Actividad biológica del ergosterol y el peróxido de ergosterol

4.4.1.1.1 Effects of temperature and pressure

The well-ordered prismatic and platy crystals observed in dome rock samples from Soufrière Hills by Horwell et al. (post review; Appendix 4) appear to be two model end-members of variable crystal habits. The presence and morphological variability of both platy cristobalite, which is interpreted to be a pseudomorph after hexagonal tridymite (Horwell et al., post review; Appendix 4), and prismatic morphologies suggests variability in the pressure-temperature regime in the dome environment, since each polymorph occupies a different P-T stability field (Figure 2.1). The presence of platy cristobalite also implies a paramorphic transformation to cristobalite. Horwell et al. (post review; Appendix 4) have tentatively attributed the presence of the two different morphologies to a quantifiable difference in crystal impurities (see Chapter 5); however, the atomic structure of a crystal is also governed by P-T conditions, so we further propose that the type of polymorph crystallising depends on the combined effects of temperature, pressure and source composition (elemental abundance).

Temperatures in volcanic domes are estimated to be at or below approximately 850 °C (e.g., Barclay et al., 1998; Murphy et al., 2000). Cristobalite is stable at temperatures above 1470

°C, and must, therefore, crystallise and persist metastably at dome temperatures. Tridymite, however, is stable at temperatures above 870 °C, approximately dome temperature. Both

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tridymite and cristobalite have been synthesised at temperatures below their stability field in the presence of a mineraliser, e.g., Al2(SO4)3 and Na2CO3 (Chao and Lu, 2002a; Stevens et al., 1997). Al³⁺ can go into 4-fold coordination with oxygen and can therefore substitute for Si⁴⁺. When such a substitution takes place, the charge imbalance is offset elsewhere in the silicate structure. This substitution in cristobalite can be stabilised by Na, Li, K, Ca and others (e.g.; Parise et al., 1994). Horwell et al. (2003b) identified trace amounts of aluminium and sodium in crystalline silica in Soufrière Hills ash, which has been confirmed for all current study locations (see Chapter 5). Therefore, we hypothesize that hexagonal tridymite may form initially when dome temperatures are near the tridymite stability field (i.e., at or above approximately 870 °C), whereas prismatic cristobalite will preferentially form below this temperature, facilitated by cation substitutions and according to Ostwald’s rule of stages (Nývlt, 1995). This temperature dependence may account for the less frequent occurrence of hexagonal crystals in dome rock, since domes have an estimated upper temperature of approximately 880 °C (e.g., Soufrière Hills; Barclay et al., 1998). The effects of temperature on lattice substitutions, the degree of crystallinity, and the stability of cristobalite are discussed in detail in Chapter 5.

The depth of the dome rock will affect the pressure of crystallisation. Pressure is known to influence the silica crystal lattice, with the most extreme case being that of amorphous silica precipitation at >1 GPa (Hemley et al., 1988); therefore, depth of crystallisation may affect the crystal habit. Considering the case of dome formation at Mount St. Helens, the most extreme pressure difference in a static system would be that of the cryptodome versus an exogenous dome. For dome rock samples, we assume an average density of 2100 kg m^-3 which is an approximation taken from the densities of the least and most dense cryptodome-derived pyroclasts from Mount St. Helens, ρ=1600 and 2300 kg m^-3 (Hoblitt and Harmon, 1993). A depth of 500 m is taken as representative of the modelled depth for cryptodome formation (emplacement and pre-eruptive; Donnadieu and Merle, 2001). The pressure on a representative packet of dome material can be determined using:

𝑃 = 𝜌𝑔ℎ Equation 4.2

where P is the magmastatic pressure, ρ is the density of the overlying rock mass, g is the acceleration due to gravity, and h is the depth below the surface of the volcano. From Equation 4.2, the crystallisation pressure for an average cryptodome sample was approximately 10.3 MPa, and the crystallisation pressure for an average exogenous dome sample approximately 2.1 MPa. Correspondingly, the pressure at the base of the Soufrière Hills dome is estimated around 7 MPa (Hicks et al., 2009). Woods et al. (2002) further show

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that that pressure is > 1 MPa throughout the dome (with pressures near the centre of the dome of 5-10 MPa), except in a narrow boundary layer of 5-10 m near the surface where pressure decreases to atmospheric pressure (0.1 MPa). Although the calculation here is a first order approximation as it does not consider the effects of dynamic fluid pressure in a closed conduit system of interconnected pore networks, these pressures are well within the stability field for both tridymite and cristobalite (low pressure silica polymorphs), and orders of magnitude below that of pressure-induced precipitation of amorphous silica. Therefore, it is unlikely that pressure directly controls the preferential formation of either polymorph.

Pressure can also influence the abundance of substitutions in a crystal lattice (e.g, aluminium in hornblende; Schmidt, 1992). Both tridymite and cristobalite have very open structures which allow impurity ions to be incorporated, possibly influencing which of these structural modifications is adopted. So, while pressure probably does not control the structural arrangement directly, the degree of substitution and subsequent polymorph could be affected.

The effects of pressure on lattice substitutions and morphology are further considered in Chapter 5.

4.4.1.1.2 Source of dissolved Si

Horwell et al. (post review; Appendix 4) propose two sources for dissolved Si: local redistribution and bulk transport. Local redistribution of silica occurs on a pore-size scale, where Si from the corrosion of amorphous SiO2 or phenocrysts is re-deposited as crystalline silica. This would occur since the necessary silica-saturation level in a vapour is lower for cristobalite deposition than that required for deposition of amorphous silica (Renders et al., 1995). Bulk transport of silica describes the flow of Si-bearing gases from parts of the magmatic system where P-T conditions favour silica corrosion to parts of the dome where conditions favour mineralisation (see Equation 4.1). No textural evidence, however, has been previously provided for either process.

The embayment of vesicle walls in samples from Santiaguito and Mount St. Helens (Figure 4.14) are interpreted to be corrosional embayments, and the result of vapour flux. The ‘rind’

shows characteristic devitrification textures and, therefore, the inner edge is thought to be the original vesicle wall (demarcated in Figure 4.14). This is supported by the presence of vapour-phase cristobalite originating at this inner vesicle wall. The possibility of mineral deposition or precipitation forming the ‘rind’ is discounted since it is comprised of plagioclase microlites and oxides (in addition to crystalline SiO2) with a similar distribution to the surrounding groundmass. Further, it is of a uniform thickness and mineral precipitation (e.g., of silica from water) would be expected to preferentially alter and precipitate silica on one surface. The possibility of a hydration halo, whereby a glassy rim has formed due to

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rehydration of the melt, is discounted for similar reasons (although, some may be expected).

The gap between the ‘rind’ and the pristine groundmass is interpreted as a ‘path of least resistance’, whereby fluid preferentially corrodes the more reactive glass than crystalline phases. The abundance of sub-micron vesicles throughout the ‘rind’ and their absence in the surrounding groundmass could further support the notion of interaction with corrosive fluids.

If so, this may be the first textural evidence supporting the potential of volcanic gases to scavenge material from degassing pathways in a dome and contribute to local and/or bulk vaporous transport of Si. Considering the equilibrium equation of Equation 4.1, which outlines both precipitation and corrosion of SiO2, higher pressures at depth would favour silica corrosion and lower pressures in the dome favour crystallisation. Contrary to this understanding, the Santiaguito sample was collected in situ from the surface of the dome.

However, it is unknown whether this parcel of magma was held at depth before moving to shallower regions, thereby texturally preserving this process. Further examples are necessary to constrain the process recorded by this texture.

4.4.1.2 Controls on devitrification

No studies were found that address the mechanism producing devitrification cristobalite in volcanic domes; however, multiple studies from the ceramics industry investigate cristobalite crystallisation from amorphous SiO2 under P-T conditions relevant to volcanic domes.

Direct crystallisation of cristobalite from glass is difficult and requires temperatures above 1000 °C (Sosman, 1965). However, studies on both natural rhyolitic and synthesised glass have shown that alkali-rich aqueous solutions increase devitrification rates (Bassett et al., 1972; Lofgren, 1971). In the presence of sodium and potassium salts, for example, α-cristobalite can form at 800 °C in as little as 2 hours (Bassett et al., 1972). Further, Na2O, Na2CO3, K2CO3, and NaCl and KCl-bearing vapours can reduce the crystallisation temperature to 700 °C (Bassett et al., 1972; De Keyser and Cypres, 1961). Both studies are within the range of temperatures expected for volcanic domes.

The silica polymorph formed is dependent on the nature and size of cations incorporated within the structure, with Na⁺ and K⁺ leading to a less dense cristobalite structure (Bassett et al., 1972). Volcanic devitrification cristobalite in the current study, as well as that by Horwell et al. (accepted; Appendix 4), is known to contain up to 1.5 wt. % Na (see Chapter 5), potentially implicating alkali substitutions in reducing the crystallisation temperature in natural settings. It is likely that Na⁺ preferentially substitutes compared to K⁺ because of its smaller ionic radius (0.96 Å versus 1.3 Å), which facilitates the coupled, electrically-neutral substitution of Al³⁺ and Na⁺ for Si⁴⁺ in the lattice (see Chapter 5). The structural substitution of Al³⁺ and Si⁴⁺ is common in rock forming minerals (Deer et al., 1996) and universally

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occurs here due to the natural abundance of Al in volcanic glass. According to Pauling’s rules on ionic coordination (Pauling, 1929), lower crystallisation temperatures will favour the larger elemental coordination number; that is, preferential substitution of Al for Si at lower temperatures.

The charge-balancing concentrations of Na⁺ (and K⁺, experimentally) preclude the formation of the tighter quartz lattice, which has a structural void of 1.4 Å, whereas cristobalite, at 1.7 Å, can spatially accommodate larger cations. These substitutions, combined with the low enthalpy of formation for cristobalite, results in the preferential devitrification of volcanic glass to cristobalite. We note that, in general, high temperature quartz has relatively greater levels of aluminium compared with low temperature quartz (e.g., Dennen et al., 1970; Perry, 1971; Scorrono, 1975); however, we hypothesize that this is directly related to increased lattice spacing at higher temperatures (Chapter 5), permitting the Si⁴⁺ substitution by Al³⁺

alongside interstitial trace cations (e.g., H and Li; Miyoshi et al., 2005).

The temperature required for rapid crystallisation (generally considered 2 hours experimentally) increases with increased cation size (e.g., Na⁺ < K⁺), which is also the order of increasing diffusion coefficients for the ions in glass (Bassett et al., 1972). Therefore, at the lower temperatures in volcanic domes than those used experimentally, the rate of devitrification (and corresponding cristobalite abundance in a sample) could be dependent on cation diffusion, with the greater diffusivity of Na⁺ limiting the incorporation of K⁺ and Ca²⁺

in volcanic cristobalite.

Rapid decompression can lead to a degree of undercooling which results in decompression-induced crystallisation. Hammer and Rutherford (2002) investigated the temporal evolution of feldspar crystallisation during isothermal decompression of equilibrated (780 °C, 220 MPa) dacitic Pinatubo pumice, and, in doing so, crystallised devitrification cristobalite (reproduced here as Figure 4.19, which has been adapted from their paper). Samples allowed to decompress to final pressures of 5-10 MPa (approximate dome pressures, as calculated above) contained particularly high levels of devitrification. Samples held at these pressures for 38 days were most effective at crystallising cristobalite, which although marked as quartz, show the indicative ‘fish-scale’ cracking and ‘feathery’ devitrification texture.

According to the present ranking system for devitrification in samples, these samples would be assigned a progression rank of ‘+++’ (e.g., Figure 4.9e and f). Williamson et al. (2010) estimated that cristobalite can form within hours to days, and these data lend credence to the rapidity with which cristobalite can form (i.e., < 38 days, and on the order of hours,.

experimentally), as well as to the value of dome residence time and extrusion rate (discussed in Section 4.3.2) in crystallising devitrification cristobalite

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Figure 4.19: Backscatter electron micrographs of depressurisation-induced crystallisation adapted from Hammer et al. (2002; Figures 3 and 5 in their study). a)-f) pressure and duration dependent crystallisation of groundmass showing ‘feathery’ textures (b and d) identified as quartz, and ‘fish-scale’ cristobalite at Pfinal 10 MPa (g) and 5 MPa (h).

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One sample in the experiments by Hammer and Rutherford (2002) also contains prismatic vapour phase cristobalite (Figure 4.19f). The sample used was pumice and assumed to have been cristobalite free; therefore, the vapour phase cristobalite is assumed to be an experimental product. Experiments were conducted in a closed system with no gas flux, thereby potentially supporting the local redistribution theory of Horwell et al. (post review;

Appendix 4) discussed above. This would require Si exsolution, sourcing excess Si from plagioclase phenocrysts, glass, and the formation of microlites; the only other crystals present. A similar study on cristobalite-rich dome rock might reveal textural evidence of the effects of re-equilibrating cristobalite, which could be used to re-consider their experimental samples and confirm the observed cristobalite as experimental.