Mineral Stability in Peralkaline Silicic Rocks: Information from Trachytes of the Menengai Volcano, Kenya

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7-2011 Mineral stability in peralkaline silicic rocks: Information from trachytes of the Menengai volcano, Kenya. Ray Macdonald Lancaster University

Boguslaw Baginski University of Warsaw

Philip T. Leat British Antarctic Survey

John C. White Eastern Kentucky University, [email protected]

P Dzierzanowski University of Warsaw

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Recommended Citation Macdonald, R., Bagiński, B., Leat, P.T., White, J.C., and Dzierżanowski, P., 2011, Mineral stability in peralkaline silicic rocks: Information from trachytes of the Menengai volcano, Kenya. Lithos, v. 125, p. 553-568. (doi: 10.1016/j.lithos.2011.03.011)

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Lithos 125 (2011) 553–568

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Mineral stability in peralkaline silicic rocks: Information from trachytes of the Menengai volcano, Kenya

R. Macdonald a,b,⁎, B. Bagiński a, P.T. Leat c, J.C. White d, P. Dzierżanowski a a IGMiP Faculty of Geology, University of Warsaw, al. Żwirki i Wigury 93, 02-089 Warszawa, Poland b Environment Centre, Lancaster University, Lancaster LA14YQ, UK c British Antarctic Survey, High Cross, Madingley Road, Cambridge CB30ET, UK d Department of Geography and Geology, Eastern Kentucky University, Richmond, KY 40475, USA article info abstract

Article history: Electron microprobe analyses are presented for phenocrysts and matrix glass in peralkaline, silica- Received 5 January 2011 oversaturated trachytes from the Menengai volcano, Kenya Rift Valley. The dominant phenocryst assemblage Accepted 10 March 2011 is alkali feldspar–hedenbergite–fayalite–titanomagnetite–apatite. Aenigmatite, amphibole and quartz occur Available online 17 March 2011 rarely in more peralkaline rocks. QUILF calculations indicate that the trachytic magmas crystallised at

temperatures of 854–870 °C, at relatively low oxidation states (ΔFMQ −1.6 to −1.7) and silica activity (aSiO2 Keywords: (Qtz) of 0.60–0.66). The new analyses are used, along with published data, to outline the distribution of the Menengai volcano main phases in the compositional spectrum of peralkaline quartz trachytes and rhyolites. There is uncertainty Peralkaline trachytes and rhyolites Mineral stability ﬁelds about the nature, or even existence, of a low-temperature zone in the alkali feldspar primary phase region, the Mineral composition equivalent of the thermal valley in the haplogranite system. Quartz phenocrysts may appear early or late in QUILF the crystallisation sequence, even in rocks of similar bulk composition, its appearance perhaps being a function of the F content of the melts. Whereas hedenbergite and fayalite show fairly systematic compositional trends with increasing host rock peralkalinity, amphibole compositions are variable, for reasons not yet understood. Aenigmatite crystallisation is at least partly controlled by oxygen fugacity and silica activity. With rare exceptions, ilmenite and titanomagnetite are incompatible phases but the factors controlling their relative stabilities are not clear. It appears that peralkaline trachyte–rhyolite sequences

1. Introduction comenditic trachytes. There is a notable scarcity globally of pantelleritic trachytes; for example, of the 421 analyses of peralkaline trachytes and Peralkaline salic rocks are common eruptive products in continental rhyolites compiled by Macdonald (1974), only 21 (5%) were pantelleritic rift valleys and oceanic islands and represent one of the highly evolved trachytes. Nineteen of the 21 were from the central Kenyan section of the components of intraplate magmatism. In some areas they form large East African Rift System. Pantelleritic trachytes have subsequently been volumes of eruptive rocks. For example, peralkaline rhyolites in the reported from other settings, e.g. ocean islands such as Terceira, Azores Ethiopian ﬂood sequences have a dense-rock-equivalent volume of (Self and Gunn, 1976) and Socorro Island, Mexico (Bohrson and Reid, 60,000 km3 (Ayalew et al., 2002) and peralkaline rhyolites in the Deccan 1997) but the greatest known volumes are still in the East African Rift Traps form between 500 and 1000 km3 (Lightfoot et al., 1987; Scaillet System. Volcanic centres in the Kenya Rift Valley where pantelleritic and Macdonald, 2006a). Scaillet and Macdonald (2006a) have pointed trachytes have been recorded include Longonot (Rogers et al., 2004), out the major potential environmental consequences of sulphur release Eburru (Clarke et al., 1990; Ren et al., 2006), Menengai (Leat et al., 1984; into the atmosphere by such large volumes of magma. Macdonald et al., 1970), Nasaken (Weaver et al., 1972), Silali

On the basis of Al2O3 and FeO* (total Fe as FeO) contents, Macdonald (Macdonald et al., 1995), and Emuruangogolak (Black et al., 1998; (1974) subdivided the rocks into four types: comendites, comenditic Weaver, 1977). trachytes, pantellerites and pantelleritic trachytes (Fig. 1). Peralkaline Peralkaline trachytes are of petrological and volcanological interest silicic suites can evolve along several liquid lines of descent, as for several reasons. They show extreme degrees of element depletion exempliﬁed in Fig. 1. The majority of pantellerites evolve from and enrichment at relatively low SiO2 contents; for example, Sr abundances can be less than 1 ppm and Zr and Nb contents up to 1700 ppm and 490 ppm, respectively, in rocks with SiO b65 wt.% (Leat ⁎ Corresponding author at: IGMiP Faculty of Geology, University of Warsaw, al. 2 Żwirki i Wigury 93, 02-089 Warszawa, Poland. Tel.: +48 22772 4953. et al., 1984; Macdonald et al., 1994; Rogers et al., 2004). The petrogenetic E-mail address: [email protected] (R. Macdonald). processes leading to these features have not yet been fully quantiﬁed.

554 R. Macdonald et al. / Lithos 125 (2011) 553–568

Fig. 1. FeO*–Al2O3 plot to show ﬁelds of peralkaline quartz-normative extrusive rocks: CT, comenditic trachytes; C, comendites; PT, pantellerites; P, pantellerites (from Macdonald, 1974). Selected suites exemplify the variability in generalised differentiation trends. Data sources: Pantelleria, Italy, pantellerite lavas and tuffs, White et al. (2009); precaldera lavas, Fantale, Ethiopia, Gibson (1972);Gedemsa,Ethiopia,Peccerillo et al. (2003);Boina, Ethiopia, Barberi et al. (1975); Longonot, Kenya, unit Lt2b, Rogers et al. (2004);Olkaria, Kenya, Macdonald et al. (1987), Eburru Trachyte Formation, Kenya, Ren et al. (2006); Menengai matrix glasses, this paper.

The chemical composition of peralkaline magmas can also affect their eruptive behaviour. Hydrous, halogen-rich rhyolites have extremely low viscosity (Baker and Vaillancourt, 1995; Dingwell et al., 1998; Gioncada and Landi, 2010), facilitating efficient degassing of magmas, relative ease of eruption, and efficient separation of residual melts from crystal mushes in magma reservoir. In contrast, the high Fe contents, especially of pantelleritic varieties, can lead to high magma densities, inhibiting crustal ascent and eruption (Di Carlo et al., 2010; Scaillet and Macdonald, 2006b). Where fractionation leads to higher-Fe residual melts, it is conceivable that the density gradient in the upper parts of zoned magma reservoirs may promote convective overturn (Leat et al., 1984). Despite their petrological and volcanological interest, there is very little published information on the mineralogy of peralkaline, especially pantelleritic, trachytes, particularly the phenocryst assemblages and compositions, information which is critical to quantifying their evolutionary paths and estimating the P–T–X conditions under which they crystallised. There is also a dearth of compositional data for matrix Fig. 2. Location of the Menengai caldera volcano within the central Kenya peralkaline glasses in peralkaline rocks in general, which has restricted attempts to province. Menengai, Longonot and Suswa are trachytic caldera volcanoes, whereas examine such features as approaches to mineral–melt exchange Eburru and Olkaria are rhyolitic dome fields. The Ndabibi and Akira Plains fields contain equilibria and how they vary with P–T–X conditions. Here we present peralkaline rhyolites and trachytes in addition to basalts. electron microprobe analytical (EMPA) data for trachytes of the Menengai caldera volcano, Kenya Rift Valley, with the specific aims of: (i) documenting phenocryst and matrix glass compositions and their (Macdonald and Bagiński, 2009; Macdonald and Scaillet, 2006). Suswa, relationship to bulk rock composition; (ii) examining in detail alkali Longonot and Menengai are trachytic caldera volcanoes, whereas feldspar–melt relationships; and (iii) outlining mineral stability fields Olkaria and Eburru are rhyolitic complexes. Menengai is composed for the whole spectrum of peralkaline silica-oversaturated rocks. largely of silica-oversaturated, peralkaline trachytes, with subordinate Menengai is particularly suitable for such a study: (1) there is a volumes of metaluminous trachytes and pantelleritic rhyolites. The clear genetic link between metaluminous and peralkaline trachytes; geology of the volcano has been described by Leat (1984, 1991).Briefly, (2) the transition from comenditic trachytes through pantelleritic a shield-building phase (K–Ar dated at 0.18±0.01 Ma) erupted lavas 3 trachytes to pantellerites is possibly unique; and (3) unusually for and interbedded fall deposits with an estimated volume of 29 km .This peralkaline trachytes, most, but unfortunately not all, eruptive units was followed by two caldera-forming events (at 29 ka and 8 ka, include pristine glassy samples. This is critically important for respectively), each accompanied by eruption of an ash-flow tuff with 3 peralkaline rocks, which can show significant chemical modifications a volume of 20–30 km . Post-caldera activity has largely been restricted on crystallisation and secondary hydration, particularly mobilisation to inside the second caldera and produced mainly lavas (at least 70 of the alkali and alkali earth elements. flows), sheet-forming fall pumice deposits and strombolian cinder cones. 2. Menengai volcano 3. Samples and analytical methods Menengai is a young (b0.2 Ma) caldera volcano in the south-central Kenya Rift Valley (Fig. 2). It forms part of the central Kenya peralkaline Leat et al. (1984) and Macdonald et al. (1994) presented major and province, a unique assemblage of peralkaline salic magmatic systems trace element data for all the major eruptive units at Menengai. They Author's personal copy

R. Macdonald et al. / Lithos 125 (2011) 553–568 555 also presented a preliminary set of mineral chemical analyses for the A2 and rims to phenocrysts in K1A34/1. FeTi-oxides, up to 0.3 mm across, second ash flow tuff but no data for other units. For this study, we mayformdiscreteequantcrystalsbutaremorecommonlyassociated have selected samples from each eruptive unit which cover the with, and occur as inclusions within, clinopyroxene. Small (normally whole-rock compositional spectrum. Table 1 is a list of analysed b0.1 mm), euhedral apatite prisms are almost invariably associated with samples. Leat et al. (1984) showed that the Menengai rocks define the FeTi-oxides or hedenbergite. Crystals in RJK6 are zoned, with a bright several, closely related, liquid lines of descent. The samples in this core on back-scattered electron (BSE) images. A sulphide phase has been study do not, therefore, represent a single magmatic lineage. found in only two samples, RJK6 and M8, where it is partially or totally Phenocryst and matrix glass compositions were determined by included in hedenbergite and FeTi-oxide phenocrysts. Amphibole occurs electron microprobe at the Inter-Institute Analytical Complex at in only one specimen in this study (K1A34/1), where it forms a reddish- IGMiP Faculty of Geology, University of Warsaw, using a Cameca SX- brown prism 0.3 mm long. Aenigmatite is not present as phenocrysts in 100 microprobe equipped with four wavelength detectors. The any of the samples analysed in this study. Colourless to reddish-brown accelerating voltage was 15 kV and the probe current was 20 nA for glass (melt) inclusions occur in alkali feldspar and clinopyroxene pyroxene, amphibole, olivine and spinel, and 15 kV and 10 nA and phenocrysts in RJK10b, M8 and H2. beam spot diameter of ~5 μm for feldspar, to reduce Na loss. For glass Judging from the phenocryst assemblages and the mutual inclusion analyses, 15 kV and 6–10 nA and a dispersed spot of ~10–20 μm were relationships, we suggest that the order of crystallisation of the main used. Fluorine in glass was analysed separately, using 15 kV and 40 nA phases was alkali feldspar−olivine−(clinopyroxene+oxides+apatite). with a dispersed spot. The standards and X-ray lines used were: Since all phases coexist in rocks with b5% phenocrysts, the temperature wollastonite for Ca, rutile for Ti, corundum for Al, synthetic Cr2O3 for interval over which they appeared must have been small. Cr, orthoclase for K, albite for Na, diopside for Si and Mg, hematite for Fe, synthetic NiO for Ni, rhodocrosite for Mn, barite for Ba, tugtupite 5. Glass and phenocryst compositions for Cl and synthetic fluor-phlogopite for F. The ‘PAP’ φ(ρZ) programme (Pouchou and Pichoir, 1991) was used for corrections. 5.1. Glass Apatite was analysed for REE, Th and U using 15 kV and 50 nA. The standards used were end-member synthetic phosphates (XP5014) for Compositional data for matrix glass and melt inclusions (in RJK10b each REE, YAG for Y, apatite for P, diopside for Ca and Si, and synthetic and H2) are given in Supplementary Table 1; average analyses are fluorphlogopite for F. In most analyses, 1 σ for the REE ranged from presented in Table 2. The glass in individual samples covers variable 0.02 to 0.06 wt.% and for Y b0.01 wt.%. Fuller details of apatite amounts of the total range in the glasses, from restricted (e.g. K1A1 determinations may be found in Macdonald et al. (2008). and K1A2) to extensive, the largest ranges being in rocks from the Representative or average analyses of phenocrysts and glass are second ash flow tuff (RJK6; Table 2 ). Macdonald et al. (1994) also presented in Tables 2–6; the full data sets are given in Supplementary recorded variable glass compositions in rocks from this unit and Tables 1 and 2. ascribed the range to mixing of melts from different layers in a compositionally zoned magma reservoir during magma withdrawal. At least 4. Petrography part of the variation in rocks showing more restricted ranges may, however, be due to incomplete mixing of melts at hand-specimen scale ThemajorityofMenengaieruptiverocksarephenocryst-poor,with which were locally heterogeneous because of proximity to different modal phenocrysts normally ranging from 0 to 6%. Exceptions are the last- phenocrysts. Furthermore, glass analyses are prone to contamination by erupted parts of the ash flow tuffs, with modal abundances up to 35%. microlites. In Table 2, glasses showing only small compositional ranges The dominant assemblage is alkali feldspar+olivine+clinopyroxene+ have been averaged; more- and less-Fe-rich averages are presented for titanomagnetite+apatite and occurs over a wide range of whole-rock RJK6, where the range is significant. Phenocrysts in this rock are mainly compositions (Table 1). Quartz occurs in samples A1 and A2 as rounded associated with the lower-Fe component. Glass (melt) inclusions in microphenocrysts up to ~150 μm across. Amphibole and aenigmatite are phenocrysts show no systematic differences to matrix glass in the same rare phenocrysts (Leat, 1983; Macdonald et al., 1994). Alkali feldspar, rock and have been included in the averaging. The similarity reinforces overwhelmingly the most abundant phenocryst (N90% of the assem- the point that the phenocrysts were close to being in equilibrium with blage), forms tabular or lath-shaped crystals, varying in size up to 5 mm. the host melts. Some crystals are partly resorbed; others are fragmented, especially in the The glass compositions range from silica-saturated (normative ash flow tuffs. Olivine occurs as colourless, partly resorbed, prismatic olivine 4.9%) to silica-oversaturated (normative quartz up to 17%) and grains up to 1 mm long, sometimes with a rim of FeTi-oxide. from metaluminous (PI 0.95, where PI is Peralkalinity Index, mol.

Clinopyroxene forms euhedral to subhedral, colourless to pale green, (Na2O+K2O)/Al2O3) to strongly peralkaline (PI ~1.7). With increasing crystals up to 1.5 mm across. A darker green variety forms phenocrysts in PI, there are overall increases in SiO2, FeO*, Na2O, F and Cl and decreases in TiO2,Al2O3,CaOandK2O(Fig. 3). The displacement of SiO2 from the main trend in samples A1 and A2 is probably a result of minor quartz Table 1 Phenocryst assemblages of Menengai rocks. fractionation. Sample K1A13 is consistently displaced from the main trends in these plots. This is probably a result of the removal of Na during Sample Unit Phenocrysts secondary hydration of the glass; raising the Na2O value from 6.48 to K1A13 PrCT af+ol+cpx+tm+ap 7.69 wt.% would increase the PI to 1.5, consistent with the overall A1 AFT1 af+qz composition of this rock. Maximum F and Cl values are 0.53 wt.% and A2 AFT1 af+cpx+qz 0.32 wt.%, respectively, broadly comparable to values in pantellerites H2 AFT2 af+ol+cpx+tm+ap M8 AFT2 af+ol+cpx+tm+ap+po from the nearby Eburru complex (Ren et al., 2006)butlowerthanthe RJK6 AFT2 (ET) af+ol+cpx+tm+ap+po highest values in the peralkaline rhyolites of the Olkaria complex RJK10b PoCT af+ol+cpx+tm+ap (FN1 wt.%; Cl 0.5 wt.%; Macdonald et al., 1987; Scaillet and Macdonald, K1A34/1 PoCT af+ol+cpx+tm+ap+ab 2001). KOR23 PoCL aphyric KOR32 PoCL af+ol+ap 5.2. Alkali feldspar PCT, pre-caldera tuff; AFT1, first ash-flow tuff; AFT2, second ash-flow tuff; ET — Engoshura Tuff (welded proximal facies of second ash-flow); PoCT, post-caldera; tuff; PoCL, post-caldera lava. af, alkali feldspar; ol, olivine; cpx, clinopyroxene; tm, The total range of composition is Or32–44, with the exception of a titanomagnetite; ap, apatite; po, pyrrhotite; ab, amphibole; qz, quartz. crystal rim in K1A34/1 (Or26) (Supplementary Table 1). The feldspars Author's personal copy

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Table 2 Average compositions of matrix glass.

Sample K1A13 A1 A2 H2 M8 RJK6 (i) RJK6 (ii) RJK10b KOR23 KOR32

n 6 7 10 15 3 3 8 20 8 5

SiO2 65.94 61.91 63.69 63.56 62.52 64.24 64.94 62.33 65.58 66.43

TiO2 0.67 0.64 0.67 0.77 0.83 0.81 0.71 0.85 0.71 0.69

Al2O3 11.81 11.34 12.38 15.26 15.51 15.36 13.38 16.48 12.18 10.19 FeO* 9.08 10.07 8.44 7.25 6.90 6.92 8.68 5.68 8.42 10.28 MnO 0.39 0.35 0.35 0.32 0.35 0.23 0.36 0.22 0.45 0.44 MgO 0.12 0.12 0.2 0.31 0.40 0.36 0.21 0.53 0.16 0.11 CaO 0.91 0.70 0.85 1.27 1.37 1.28 0.97 1.55 0.96 0.68

Na2O 6.48 8.82 8.75 6.80 7.30 8.03 7.68 5.81 7.53 7.16

K2O 4.68 4.21 4.52 5.11 5.41 5.11 4.79 5.70 4.81 4.48

P2O5 0.03 0.04 0.08 0.10 0.16 0.13 0.08 0.17 0.07 0.05 F 0.40 0.53 0.31 0.17 0.13 0.14 0.28 0.07 0.25 0.35 Cl 0.26 0.32 0.19 0.10 0.08 0.08 0.18 0.05 0.16 0.20 Sum 100.76 99.05 100.43 101.03 100.96 102.69 102.26 99.43 101.27 101.06 O=F,Cl 0.23 0.29 0.17 0.09 0.07 0.08 0.16 0.04 0.15 0.20 Total 100.53 98.76 100.26 100.94 100.89 102.61 102.10 99.39 101.12 100.86 PI 1.33 1.68 1.56 1.09 1.15 1.22 1.33 0.95 1.44 1.63

RJK (i) and RJK (ii) are the less- (FeO* b7.5 wt.%) and more (N7.5 wt.%) -Fe rich varieties, respectively. FeO*, total Fe as Fe2+. n, number of point analyses; fewer for F (3–6).

PI, Peralkalinity Index (mol. (Na2O+K2O)/Al2O3). thus span the anorthoclase–sanidine boundary. We discuss the the Gedemsa volcano, Ethiopia. At Menengai, the Fe contents balance relationship between feldspar and whole-rock compositions below; Al deficiency in the feldspar and it may be that the increasing whole- we note here that there is an overall decrease in Or content as host rock deficiency in Al is being matched by that in the feldspar. Barium glass PI increases. The An content varies from 0 to 10 mol% and levels are significant only in the metaluminous rock RJK10b (BaO decreases with increasing rock peralkalinity. Iron contents (as Fe2O3*) 0.38 wt.%). are less than 0.5 wt.% when PI is less than 1.3 but then increase to a The within-crystal and within-rock compositional ranges vary maximum of 3.5 wt.% with increasing PI, the higher values normally from Or 0.7 to 11.5 mol%, the range generally being smaller in more being found in crystal rims. Values exceeding 1.5 wt.% are unusual in peralkaline rocks. Variation is generally not related to position within alkali feldspars from peralkaline rocks, although Peccerillo et al. crystals; however, some grains have slightly more potassic cores (2003) recorded 2.26 wt.% in a phenocryst rim from a pantellerite of grading to rims slightly richer in Na and Ca (e.g. RJK10b, H2 and A2;

Table 3 Average analyses of clinopyroxene phenocrysts.

Sample K1A13 A2 H2 M8 RJK6 RJK10b K1A34/1 K1A34/1

Lower Fe Higher Fe

n 4 16 37 7 10 38 29 36

SiO2 48.94 38.92 50.34 50.60 50.30 50.93 50.52 49.96

TiO2 0.64 9.03 0.63 0.64 0.66 0.71 0.50 0.48

Al2O3 0.35 1.57 0.52 0.60 0.54 0.86 0.53 0.44

Cr2O3 0.01 0.01 0.01 0.03 0.00 0.01 0.01 0.01

V2O3 0.02 0.04 0.02 0.01 0.01 0.01 0.01 0.01 FeO* 26.40 41.47 19.30 17.86 19.38 14.62 18.07 20.58 NiO 0.02 0.01 0.01 0.01 0.02 0.02 0.02 0.01 MnO 1.41 1.14 1.16 1.14 1.18 0.87 1.00 1.08 MgO 2.76 0.54 7.18 8.13 7.16 10.41 7.67 5.86 CaO 19.30 1.27 20.60 20.54 20.33 20.62 20.34 20.24

Na2O 0.78 6.59 0.59 0.58 0.61 0.50 0.99 0.94 Total 100.63 100.57 100.37 100.16 100.19 99.56 99.67 99.61

Formulae based on 6 oxygens Si 1.968 1.586 1.964 1.965 1.967 1.957 1.969 1.974 Ti 0.019 0.277 0.018 0.019 0.019 0.020 0.015 0.014 Al 0.017 0.075 0.024 0.028 0.025 0.039 0.024 0.021 Cr 0.001 0.000 0.000 0.001 0.000 0.000 0.000 0.000 V 0.000 0.001 0.001 0.000 0.000 0.000 0.001 0.002 Fe3+ 0.069 0.719 0.055 0.048 0.049 0.043 0.083 0.076 Fe2+ 0.819 0.694 0.575 0.532 0.585 0.427 0.506 0.606 Ni 0.001 0.000 0.000 0.000 0.001 0.000 0.001 0.000 Mn 0.048 0.039 0.038 0.038 0.039 0.028 0.033 0.036 Mg 0.166 0.033 0.418 0.471 0.417 0.596 0.445 0.345 Ca 0.832 0.055 0.861 0.855 0.851 0.849 0.849 0.857 Na 0.061 0.520 0.045 0.044 0.046 0.038 0.075 0.072 Cation sum 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 Ca 45.8 7.1 46.5 46.0 45.9 45.3 47.3 47.4 Mg 9.1 4.2 22.5 25.3 22.5 31.8 24.6 19.0 Fe 45.1 88.7 31.0 28.7 31.6 22.8 28.1 33.6 Mg-number 16 2 40 45 40 56 43 34

FeO*, total Fe as Fe2+. Mg-number, 100 Mg/(Mg+Fe2+). Fe3+ and Fe2+ from stoichiometry. Author's personal copy

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Table 4 Average analyses of olivine phenocrysts.

Sample KIA13 H2 M8 RJK6 K1A34/1 KOR32

n 41381243

SiO2 29.47 30.91 31.60 31.35 31.63 29.25

TiO2 0.09 0.06 0.09 0.05 0.04 0.07

Cr2O3 0.01 0.01 0.01 0.02 0.01 0.01 FeO* 65.59 58.70 56.10 56.32 55.60 65.66 MnO 3.79 3.59 3.44 3.53 3.60 3.80 NiO 0.02 0.01 0.02 b.d. 0.02 0.01 MgO 1.22 6.64 9.06 8.07 8.59 1.01 CaO 0.55 0.65 0.57 0.65 0.47 0.41 Total 100.71 100.58 100.87 99.99 99.95 100.22

Formulae based on 4 oxygens Si 0.986 0.992 0.993 0.988 1.003 0.985 Ti 0.002 0.002 0.002 0.001 0.001 0.002 Cr 0.000 0.000 0.000 0.000 0.000 0.000 Fe2+ 1.834 1.575 1.474 1.500 1.474 1.848 Mn 0.108 0.098 0.092 0.095 0.096 0.108 Ni 0.000 0.000 0.001 0.000 0.000 0.000 Mg 0.061 0.318 0.424 0.383 0.406 0.051 Ca 0.020 0.022 0.019 0.022 0.016 0.015 Cation sum 3.01 3.01 3.00 3.00 3.00 3.01 Fo 3.0 16.0 21.3 19.4 20.5 2.5 Fa 91.6 79.1 74.1 75.8 74.6 92.1 Tp 5.4 4.9 4.6 4.8 4.9 5.4 Mg-number 3.2 16.8 22.3 20.3 21.6 2.7

FeO*, total Fe as Fe2+. Mg-number, 100 Mg/(Mg+Fe2+). b.d., below detection.

Supplementary Table 1). In a study of peralkaline rhyolites from Gran with fractionation curves being channelled more quickly towards it, and Canaria and by analogy with studies of zoning in plagioclase, Troll and trends towards increasingly sodic compositions. Feldspar-rock data from Schmincke (2002) considered ranges of the order 3–4 mol% to lie Macdonald et al. (1970) were used by Roux and Varet (1975) to construct within the compositional range of local disequilibrium and/or small a fractionation curve for Menengai; it was clear that the orientation of the T–P changes during crystal growth. On that basis, it would appear that curve, showing strong Na-enrichment, was very different to any other the Menengai feldspars, especially in the less peralkaline rocks, did peralkaline suite modelled by them. not always reach complete equilibrium with melt. The new glass-feldspar data from Menengai (Fig. 5b) conﬁrm the The nature of alkali feldspar–liquid relationships in peralkaline strong Na-enrichment in the melts. The feldspars are in all cases more silica-oversaturated rocks was keenly debated by, inter alia, Bailey potassic than the coexisting melt, demonstrating the orthoclase effect (1974), Bailey and Schairer (1964), Carmichael and MacKenzie (1963), of Bailey and Schairer (1964), such that feldspar crystallisation Macdonald et al. (1970), Nicholls and Carmichael (1969), Roux and generates more sodic residual melts. The tie-line of one of the most Varet (1975) and Thompson and MacKenzie (1967). More recently the evolved rocks, KOR32, is at an angle to the others. Whereas it might be topic has been relatively neglected, at least partly due to a paucity of argued that this represents movement of residual melts towards a relevant studies of natural peralkaline sequences. A central issue in the low-temperature zone, the rock still demonstrates the orthoclase debate is the form in increasingly peralkaline compositions of the low- effect. Further decreases in Na/K ratio of residual melts would require temperature zone identiﬁed in the haplogranite system, Quartz– that the rocks precipitate more sodic feldspar (AbN80). This would Orthoclase–Albite, by Tuttle and Bowen (1958).Wenowusealkali further require that the feldspars move up-temperature from the feldspar–melt (glass) relationships in the Menengai rocks to comment minimum between the alkali feldspar solid solution loops (Or30 at on the form of the low-temperature zone within the alkali feldspar pH2O=1 kbar pressure (Bowen and Tuttle, 1950); Or35 at 1 atm. primary phase region in peralkaline silica-oversaturated systems. anhydrous (Schairer, 1950)). We know of no case where such strongly It is important that projections illustrating the low-temperature zone sodic feldspars have been found in a peralkaline rhyolite (c.f. Bailey, must simultaneously show variations in melt Na/K ratios, peralkalinity 1974). and degree of silica-oversaturation (Bailey and Macdonald, 1969; Roux Fig. 5b compares alkali feldspar–whole-rock (or glass) composi- and Varet, 1975). Fig. 4,afterRoux and Varet (1975),plotsparameter tions for rocks from Menengai, Eburru and Pantelleria. On the FeO*–

Y (as a measure of silica-enrichment) against parameter Z (peralkalinity). Al2O3 plot (Fig. 1), these suites have increasingly lower Al2O3 contents Any series which evolves dominantly by alkali feldspar fractionation will at the same FeO* content. The Eburru and Pantelleria data are shown as form a straight line passing through the origin (feldspar composition) on fields enclosing feldspar–rock (or glass) tie-lines. The field of Eburru this plot; generalised trends from the Boina volcano and Pantelleria pantellerites is at a slight angle to the Pantelleria rocks, whilst the exemplify this behaviour. The trend of the Menengai glasses is quasi- pantelleritic trachytes from Menengai and Eburru are at an angle to linear, the scatter reflecting the fact that in detail the rocks form several both the Eburru and Pantelleria fields. The compositional range of the lineages (Leat et al., 1984). The X (measuring K/Na ratio)–Yplot alkali feldspars is relatively narrow, Or35–41. It appears, therefore, that (measuring silica-excess with respect to alkali feldspar) (Fig. 5a) was feldspar–rock relationships, especially the degree of silica-enrichment, used by Roux and Varet (1975) to determine the position of the low- vary with the bulk composition of the system and that, at least in the temperature zone for a number of natural rock series. The trend of the total compositional range of these samples, there is no migration of whole-rock analyses and the orientation of the rock-feldspar tie-lines residual melts towards a single low-temperature zone. were taken to define the inferred fractionation curve for each series. The Studies of other suites may help resolve the nature, or indeed the curves taken together define a broad low-temperature zone (thermal existence, of the low-temperature zone in peralkaline systems. It is valley). Roux and Varet (1975) suggested that as bulk compositions critical that, to maintain the magmatic abundances of the alkalis in the become more peralkaline, the low-temperature zone becomes steeper, matrix material and to avoid distortions in the projections (Figs. 4 and 5), Author's personal copy

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Table 5 5.3. Quartz–feldspar relationships Representative analyses of titanomagnetite phenocrysts.

Number 1 2 3 4 5678 Although quartz phenocrysts are absent from all but the most evolved Menengai rocks (A1 and A2), we comment here on quartz– Sample K1A13 H2 H2 M8 M8 RJK6 RJK6 RJK10b feldspar relationships in peralkaline quartz trachytes and rhyolites, SiO2 0.07 0.08 0.03 0.06 0.07 0.22 0.02 0.03 partly to help establish the position of the quartz–feldspar cotectic in TiO 26.32 27.07 27.27 27.05 28.08 27.18 28.07 26.57 2 such compositions, and partly to assess the suggestion that quartz is a Al2O3 0.25 0.59 0.63 0.84 0.83 0.70 0.59 1.31

V2O3 0.15 0.10 0.14 0.06 0.14 0.07 0.13 0.07 late-crystallising phase, perhaps restricted in some cases to formation Cr2O3 b.d. 0.02 0.02 b.d. b.d. b.d. b.d. b.d. during ascent towards the surface during eruption (Di Carlo et al., 2010). FeO* 70.72 68.48 67.32 67.97 67.33 67.33 67.64 67.18 This has clear implications for its potential status as a fractionating MnO 1.52 1.82 1.71 1.81 1.69 1.90 1.90 1.34 phase. MgO 0.16 0.68 0.62 0.98 1.02 0.82 0.62 1.61 ﬁ ZnO 0.03 0.16 b.d. 0.08 0.14 b.d. 0.08 b.d. In Fig. 6 and certain subsequent gures, we plot analyses in the Total 99.22 99.00 97.74 98.85 99.30 98.22 99.05 98.11 FeO*–Al2O3 diagram, which is a convenient way of expressing the major Xusp 73.5 75.9 77.6 76.0 78.8 77.2 78.7 75.5 element compositions of peralkaline silicic rocks. Macdonald (1974)

1, subrounded, 80×55 μm; 2, euhedral, 150×150 μm; 3, bladed, 120×60 μm, in showed that the diagram can be contoured for PI (involving Al2O3,Na2O μ μ μ pyroxene; 4, anhedral, 50×22 m; 5, euhedral, 170×150 m; 6, subhedral, 85×60 m; and K2O contents) and degree of silica-oversaturation (expressed as 7, subhedral, 65×60 μm; 8, core, subhedral, 200×145 μm, in pyroxene. b.d., below normative quartz). Since SiO2,Al2O3, FeO*, Na2OandK2OmakeupN90% detection. of the composition of the rocks, plotted positions give a good representation of overall bulk composition. In these plots, we have such studies are made on non-hydrated glassy rocks or on crystalline focused on data from glassy rocks, thereby ensuring that the rocks where it can convincingly be demonstrated that there has been no phenocrysts are true phenocrysts and have not been compositionally post-emplacement alkali migration. modiﬁed during late-stage crystallisation or subsolidus events.

Table 6 Analyses of apatite (micro)phenocrysts.

Number 1 2 3 4 5 6 7 8

Sample K1A13 H2 M8 RJK6 RJK6 RJK10b K1A34/1 K1A34/1

Host Glass Oxide Glass Glass Glass Pyroxene Pyroxene

Size (μm) 120×80 200×100 45×10 40×40 55×50 135×110 Core 70×70 rim

P2O5 38.10 40.27 39.65 39.95 40.31 40.03 35.25 39.52

SiO2 1.83 0.73 0.33 0.42 0.46 0.76 3.13 1.04

Y2O3 0.36 0.17 0.10 0.09 0.13 0.18 0.58 0.22

La2O3 1.15 0.54 0.16 0.23 0.40 0.34 2.02 0.75

Ce2O3 2.32 0.91 0.49 0.66 0.69 0.74 3.72 1.48

Pr2O3 0.30 0.06 b.d. 0.09 0.27 0.02 0.44 0.38

Nd2O3 1.04 0.47 0.33 0.22 0.32 0.31 1.88 0.78

Sm2O3 0.15 0.06 0.11 0.07 0.06 0.03 0.11 0.18

Gd2O3 0.19 0.20 0.26 b.d. 0.04 0.23 0.28 0.20

Dy2O3 b.d. 0.08 b.d. b.d. 0.06 0.04 0.02 b.d.

Ho2O3 0.25 b.d. 0.10 0.12 b.d. 0.15 0.13 0.31

Er2O3 b.d. 0.07 b.d. b.d. 0.14 b.d. 0.05 b.d.

Yb2O3 0.08 b.d. 0.02 b.d. b.d. b.d. b.d. b.d. CaO 51.72 52.81 52.98 53.51 53.15 54.04 49.61 53.73 MnO 0.15 0.19 0.23 0.17 0.19 0.16 0.12 0.21 FeO* 0.57 0.72 1.50 0.72 0.50 0.52 0.59 0.59 F 4.38 3.62 3.35 5.12 3.80 4.00 4.60 3.14 Cl 0.03 0.07 0.07 0.06 0.08 0.12 0.04 0.02 Sum 102.62 100.98 99.69 101.44 100.60 101.67 102.55 102.54 O≡F,Cl 1.85 1.55 1.43 2.17 1.62 1.71 1.95 1.32 Total 100.77 99.43 98.26 99.27 98.98 99.96 100.60 101.22

Formulae based on 26 oxygens Ca 9.574 9.729 9.907 9.780 9.807 9.889 9.399 9.895 Fe 0.082 0.104 0.219 0.103 0.072 0.074 0.087 0.085 Mn 0.022 0.028 0.034 0.025 0.028 0.023 0.018 0.030 La 0.073 0.034 0.010 0.014 0.025 0.021 0.131 0.048 Ce 0.147 0.057 0.031 0.041 0.044 0.046 0.241 0.093 Pr 0.019 0.004 0.000 0.006 0.017 0.001 0.028 0.024 Nd 0.064 0.029 0.021 0.013 0.020 0.019 0.118 0.048 Sm 0.009 0.004 0.007 0.004 0.004 0.002 0.007 0.011 Gd 0.011 0.011 0.015 0.000 0.002 0.013 0.017 0.011 Dy 0.000 0.005 0.000 0.000 0.003 0.002 0.001 0.000 Ho 0.014 0.000 0.006 0.007 0.000 0.008 0.007 0.017 Er 0.000 0.004 0.001 0.001 0.008 0.000 0.003 0.000 Yb 0.004 0.000 0.001 0.000 0.000 0.000 0.000 0.000 Yb 0.033 0.016 0.009 0.008 0.012 0.016 0.054 0.020 Sum M 10.05 10.02 10.26 10.00 10.04 10.12 10.11 10.28 P 5.571 5.860 5.856 5.767 5.875 5.785 5.275 5.748 Si 0.316 0.125 0.058 0.072 0.079 0.129 0.553 0.178 Sum Z 5.89 5.98 5.91 5.84 5.95 5.91 5.83 5.93 Cation sum 15.9 16.0 16.2 15.8 16.0 16.0 15.9 16.2

2+ Al2O3 and SrO below detection in all samples. FeO*, total Fe as Fe . b.d., below detection. Author's personal copy

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Fig. 3. Composition of matrix glass and melt inclusions from Menengai plotted against Peralkalinity Index (mol. (Na2O+K2O)/Al2O3). Data from Table 2. Average compositions used, except for RJK6, where the averages of more Fe-rich and less Fe-rich varieties are connected by tie-lines. Samples A1 and A2 are arrowed.

Peralkaline trachytes and rhyolites are plotted in Fig. 6, distinguishing those with, and without, quartz phenocrysts. Two fields can be distinguished; one where quartz is absent and one where it may or may not be present. The boundary between the fields lies very close, except at the most Fe-rich end, to the 20% normative quartz contour drawn by Macdonald (1974), indicating that the appearance of quartz is at least partly composition-controlled. The question arises as to why only some rocks in field B are quartz- phyric. In their experimental studies of three comendites from the Olkaria complex and a pantellerite from the Eburru complex, Scaillet and Macdonald (2001, 2003, 2006b) found that the alkali feldspar-in and quartz-in crystallisation curves appeared close together and that the phases coexisted throughout the full crystallisation history. Quartz phenocrysts are ubiquitous in the natural rocks at Olkaria, sometimes forming graphic intergrowths with alkali feldspar (Macdonald et al., 1987; Marshall et al., 2009) and mass-balance modelling of compositional variations in the suite showed that quartz was an important Fig. 4. Z (=((K+Na)/Al)−1) v. Y (=(Si/3Al)−1) plot (after Roux and Varet, 1975)to component of the fractionating assemblage (Macdonald et al., 1987; show that the Menengai glass compositions are broadly compatible with alkali feldspar fractionation within a group of related magmatic lineages. Generalised Boina and Marshall et al., 2009). On the other hand, quartz phenocrysts are not Pantelleria trends shown for comparison (from Roux and Varet, 1975). abundant in the Pantellerian peralkaline rhyolites and, where present, Author's personal copy

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comparable peralkalinity (F~0.2 wt.%; Mahood and Stimac, 1990; White

et al., 2009;H2O4–4.5 wt.%; Di Carlo et al., 2010; Gioncada and Landi, 2010). Since the effect of both volatiles in the Quartz–Orthoclase–Albite system is to increase the size of the quartz ﬁeld relative to that of alkali feldspar (Manning, 1981; Manning et al., 1980), and if we assume that Cl does not have the same effect as F, we suggest that the Kenyan rocks have reached the natural quartz–feldspar cotectic at a rather earlier stage of crystallisation than the Pantellerian rhyolites. Some support for this suggestion comes from other peralkaline sequences. The Eburru rocks have F in excess of Cl and are commonly quartz-phyric, the quartz phenocrysts often occurring in graphic intergrowth with alkali feldspar (Ren et al., 2006). In Mayor Island rhyolites, many of which lack quartz phenocrysts, Cl exceeds F, although Cl/F ratios are lower than in Pantelleria pantellerites (Weaver et al., 1990). Similarly, the pantellerites of the Pico Alto volcano, Terceira, lack quartz phenocrysts and, although data are scarce, apparently have ClNF(Mungall and Martin, 1995, 1996). The Emory Peak Rhyolite, a high-silica comendite which crops out in the Chisos Mountains, west Texas, has a phenocryst assemblage dominated by sanidine and quartz, with accessory magnetite, biotite, monazite and zircon (White et al., 2006). The biotite is nearly pure annite and high in F (2.2 wt.%). This, combined with presence of ﬂuorine-rich arfvedsonite in the groundmass, led White et al. (2006) to suggest that the unit had high F and high F/Cl.

5.4. Clinopyroxene

Clinopyroxene phenocrysts are present in seven of the nine porphyritic samples (Table 1). The majority are hedenbergite (average compositions are shown in Fig. 7), although the phenocrysts in RJK10b are augite (Table 3). The degree of zonation is variable but generally small, ranging from b2FsinA2to~5FsinRJK10bandH2 (Supplementary Table 2). Sample K1A34/1 is unusual in containing

two pyroxene populations, one averaging Ca47.3Mg24.6 Fe28.1, the other Ca47.4 Mg19.0Fe33.6 (Table 3). Where present, zoning is usually reverse. Na contents are low (≤0.14 a.p.f.u.) and increase with the Fe content of the clinopyroxene and the peralkalinity of the host. With the exception of some point analyses in RJK10b, the pyroxenes have Na NAl. Hedenbergite crystallisation would, therefore, tend to reduce the peralkalinity of residual melts. Aegirine-augite occurs as phenocryst rims in K1A34/1 (Ti 0.02 a.p.f.u.; Al 0.02 a.p.f.u.) and as discrete phenocrysts in A2, distinctly richer in Ti and Al (Ti 0.27–0.28 a.p.f.u.; Al 0.06–0.09 a.p.f.u.).

Clinopyroxene–melt Mg–Fe exchange distribution coefficients (KD) values are in the narrow range 0.12–0.13 and are thus independent of host-rock peralkalinity. The values are comparable to those (0.16 and 0.15) determined on hedenbergite from Pantellerian peralkaline trachytes and rhyolites by Carmichael (1962) and to those (0.12–0.15, Fig. 5. Plots of X (=K/(Na+K)) against Y (=(Si/3Al) −1), after Roux and Varet (1975). with one value of 0.29) of Mahood and Stimac (1990). Di Carlo et al. (a) Boina and Pantelleria exemplify magmatic suites used by Roux and Varet to identify (2010) reported a value of 0.14 from phenocrysts in a Pantellerian a low-temperature zone within the alkali feldspar primary phase region for peralkaline trachytes and rhyolites. Their Menengai trend is at a high angle to those (and all other) pantellerite and of 0.16 for hedenbergite synthesised from the same fi suites. (b) The tie-lines connect alkali feldspar and glass compositions in Menengai rock. The KD values are lower than those determined for ma cand trachytes; the tie-line for a pantelleritic trachyte from Pantelleria is also shown intermediate magmas, for which the equilibrium value is 0.27±0.03 (arrowed). The fields enclose the range of alkali feldspar–glass tie-lines for pantellerites (Putirka, 1999; Putirka et al., 2003) and seems to reflect the tendency in – from Pantelleria and alkali feldspar whole-rock tie-lines for Eburru. Note the different peralkaline systems for Fe to be preferentially partitioned into the melt. overall orientations of the tie-lines in the three groups. Data sources: Pantelleria, – Carmichael (1962), Mahood and Stimac (1990); Eburru, Bailey (1974), Ren et al. It is clear from an FeO* Al2O3 plot (Fig. 8a) that clinopyroxene can (2006); Menengai, this paper and Bailey et al. (1974). crystallise over the whole range of peralkaline oversaturated compositions, although it is uncommon in rocks with FeO*N9 wt.%. Over the range, the great majority are hedenbergites or sodian hedenbergites. Aegirine-augite phenocrysts (as opposed to the rims of commonly occur as microphenocrysts (Avanzinelli et al., 2004; Di zoned crystals) occur in some pantellerites (Pantelleria, Carmichael, Carlo et al., 2010; Mahood and Stimac, 1990; White et al., 2005, 2009). 1962; White et al., 2005; Mayor Island, Ewart et al., 1968; Socorro, White et al. (2009) were able to model compositional variations in the Bryan, 1976; Eburru, Ren et al., 2006) and in a pantelleritic trachyte pantellerites without quartz in the fractionating assemblages. from Menengai, all with PIN1.5. The occurrence of aegirine-augite is A possible explanation for the earlier appearance of quartz in the not related in any simple way to a compositional parameter, such as PI

Kenyan rocks is that they are higher in F (N1wt.%)andH2O(4–6wt.%; or FeO* content. If fO2 is a controlling factor, its effect must be very Scaillet and Macdonald, 2001, 2006b) than those from Pantelleria of subtle because peralkaline oversaturated volcanic suites evolve at, or Author's personal copy

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traverses (up to 16 points) across two crystals showed that the variation does not vary systematically from core to rim.

Olivine–melt Mg–Fe exchange distribution coefﬁcients (KD) values for Menengai olivines are plotted in Fig. 9, along with the data from other localities for which coexisting glass and olivine data are available. The majority of values are in the range 0.3–0.4. However, two Menengai olivines and one experimental olivine from an Olkaria rhyolite have values N0.5. Determining whether these higher values are real or at least partly a result of errors associated with determining the low MgO levels (commonly b0.2 wt.%) in glass will require further work on peralkaline suites. Like clinopyroxene, olivine can crystallise from magmas over the full compositional spectrum but is relatively uncommon with rocks with FeO*N9 wt.% (Fig. 8b).

5.6. FeTi-oxides Fig. 6. FeO*–Al2O3 plot showing fields of quartz-phyric and quartz-aphyric peralkaline silicic volcanic rocks. The 20% normative quartz contour is from Macdonald (1974). Data sources for this and Figs. 8, 10–12: Pantelleria, Avanzinelli et al. (2004), Only titanomagnetite has been found as oxide phenocrysts in the Carmichael (1962), Di Carlo et al. (2010), Mahood and Stimac (1990), Villari (1974), Menengai trachytes. It shows small intra-sample variation (b2% Xusp) White et al. (2005, 2009); Azores, Schmincke and Weibel (1972), White et al. (1979); and small overall variation (Xusp 75.3–78.7, with one value of 73.5) Thorsmörk, Jørgensen (1980); Boina, Barberi et al. (1975); Fantale, Gibson (1972), (Table 5; Supplementary Table 2). Webster et al. (1993); Gedemsa, Peccerillo et al. (2003); Menengai, this paper; Eburru, The distribution of FeTi-oxides in comenditic and pantelleritic Ren et al. (2006); Olkaria (Naivasha), Macdonald et al. (1987), Marshall et al. (2009), Nicholls and Carmichael (1969); Kane Springs Wash, Novak and Mahood (1986); rocks is shown in Fig. 10, distinguishing the spinel and rhombohedral McDermitt caldera complex, Wallace et al. (1980); Sierra La Primavera, Mahood phases where data allow. The two phases occupy broadly similar fields, (1981); Socorro, Bohrson and Reid (1997), Bryan (1976); Mayor Island, Ewart et al. covering the spectrum of whole-rock compositions. Coexisting oxides (1968), Nicholls and Carmichael (1969); D'Entrecasteaux Islands, Smith (1976). are relatively rare in these rocks (Nicholls and Carmichael, 1969) but have been recorded in pantelleritic trachytes of Eburru by Ren et al. just below, FMQ. Aegirine is never found as a phenocryst phase, (2006) and in Pantellerian trachytes and rhyolites (Carmichael, 1962, although it is common as microcrysts in matrix glass. The relative scarcity of Na-rich pyroxenes severely restricts their potential to influence fractionation trends in the host magmas.

5.5. Olivine

Olivine phenocrysts are present in Menengai rocks with matrix glass PI up to 1.52. They have not been found in rocks containing aegirine-augite, aenigmatite or amphibole phenocrysts. Macdonald et al. (1994) reported olivine Fo35 in RJK10b; the range in this study is Fo23–2, although there are no analyses between Fo16 and Fo3 (Table 4; Supplementary Table 2). The Fe content increases with host-glass peralkalinity, as reported for various peralkaline suites by Mahood and Stimac (1990), Bohrson and Reid (1997), Peccerillo et al. (2003), White et al. (2005, 2009) and Marshall et al. (2009). The high contents of MnO (≤4.3 wt.%) are characteristic of peralkaline rhyolites (Carmichael, 1962). Levels of CaO reach 0.8 wt.% (0.03 a.p.f.u.); they do not vary systematically with Fo. The olivine is essentially unzoned, the range in

Fo being b1, except in K1A34/1, where the range is Fo20.2–24.5. Detailed

Fig. 8. FeO*–Al2O3 plots to show the occurrence in peralkaline trachytes and rhyolites of Fig. 7. Average composition of clinopyroxene and olivine phenocrysts in Menengai (a) clinopyroxene and (b) olivine phenocrysts. The dashed line in (a) is the 1.5 rocks plotted in the Di–He–En–Fs system. Tie-lines connect coexisting mineral pairs. Peralkalinity Index contour (higher PI at higher FeO* values; from Macdonald, 1974). The shaded areas are ﬁelds in which pyroxenes are not normally found (from Deer et al., Aegirine-augite phenocrysts are restricted to rocks with PI N1.5 but are uncommon 1997, Fig. 1). relative to hedenbergite. Data sources as in Fig. 6. Author's personal copy

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Fig. 11. (Ca +Al) plotted against (Na+Si) to demonstrate substitution among the Fig. 9. Olivine–melt Mg–Fe exchange distribution coefﬁcients (KD) plotted against peralkalinity of the host glass for various suites. Data sources: Pantelleria, Mahood and minor components in aenigmatite phenocrysts. Data sources as in Fig. 6. Stimac (1990); Olkaria, Marshall et al. (2009) and Scaillet and Macdonald (2003, experimental); Menengai; this paper. Macdonald et al. (1987, 1994), Mahood and Stimac (1990), Nicholls 1967; Mahood and Stimac, 1990; White et al., 2005, 2009), where and Carmichael (1969), Peccerillo et al. (2003), Ren et al. (2006) and the relative proportion of ilmenite and Xusp in the titanomagnetite White et al. (2005). Nicholls and Carmichael (1969) drew attention increase with host-rock peralkalinity. It is seldom clear what stabilises to the restricted compositional variation; Ti/(Ti+Fe), for example, one oxide over the other. White et al. (2005) suggested that, in the varies only from 0.14 to 0.17, despite signiﬁcant variation in whole-

Pantellerian suite, magnetite crystallised at the highest temperatures rock composition. There are, however, signiﬁcant ranges in Al2O3 and and at relatively high oxygen fugacities, whereas ilmenite crystallised CaO contents and compositional variation can be expressed by the from all but the most peralkaline, lowest-temperature, melts. Mungall coupled substitution SiIV +NaVIII ↔AlIV +CaVIII (Kunzmann, 1999) and Martin (1995) distinguished two evolutionary modes among the (Fig. 11). Overall, there is no simple relationship between whole-rock peralkaline rhyolites of Terceira, Azores. The pantelleritic suite of Pico peralkalinity and the degree of (Al+Ca) substitution. However,

Alto evolved at low fO2, with ilmenite as the oxide phase, whereas the Mahood and Stimac (1990) noted that in Pantellerian pantellerites, comenditic magmas of Santa Barbara crystallised a spinel phase, aenigmatite phenocrysts contain more Na and less Al and Ca as whole- related to higher fO2. In the Olkaria rhyolites, Kenya, ilmenite and rock PI increases. titanomagnetite occur separately in rocks over the same PI range and With the exception of some comendites close to the comendite– both coexist with varying combinations of fayalite, hedenbergite, pantellerite boundary, crystallisation of aenigmatite is restricted to amphibole and biotite. Marshall et al. (2009) noted that in the complex pantellerites, the lower limit of PI being ~1.25 (Fig. 12). In their

P–T–X–fO2 space occupied by the Olkaria rhyolite magmas, the two experimental study of a Pantellerian pantellerite, Di Carlo et al. (2010) oxides have separate stability fields. As at Olkaria, the relative stability also found that near-liquidus aenigmatite required a melt peralkalinity of the two oxides in the complete data set is not simply related to the higher than 1.2. On the basis of a generally antipathetic relationship nature of the coexisting mafic phases; ilmenite and titanomagnetite between aenigmatite and FeTi-oxides, Nicholls and Carmichael (1969) both coexist with various combinations of all the other mafic phases. proposed that aenigmatite crystallisation is stabilised by the reaction of pyroxene and ilmenite with Na-rich melt. They recognised a “no-oxide

5.7. Aenigmatite field” in T–fO2 space in which aenigmatite is stable but the oxides are not. An analysis of aenigmatite phenocrysts in the second Menengai ash There is, in some suites, textural evidence of such reactions. For flow tuff was presented by Macdonald et al. (1994). Analyses in example, Ren et al. (2006) reported ilmenite inclusions in aenigmatite peralkaline rhyolites may be found in Avanzinelli et al. (2004), Bryan phenocrysts in an Eburru pantellerite. Other suites do not show the (1976), Carmichael (1962), Di Carlo et al. (2010), Ewart et al. (1968), antipathetic relationship between aenigmatite and ilmenite; White et al. (2005) recognised five different phenocryst assemblages in

Fig. 10. FeO*–Al2O3 plot to show the crystallisation ranges of FeTi-oxide phenocrysts. Fig. 12. FeO*–Al2O3 plots of (a) aenigmatite and (b) amphibole phenocryst occurrences Data sources as in Fig. 6. in peralkaline suites. Data sources as in Fig. 6. Author's personal copy

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Pantellerian pantellerites, two of which included aenigmatite and which often have aenigmatite+ilmenite assemblages, lie largely in the ilmenite. They proposed that the two phases were in equilibrium over field where both would be stable. Magnetite-bearing samples from a large compositional range (~67–72 wt.% SiO2 and PI 1.61–1.75) and Menengai, with lower values of ΔFMQ as noted earlier, lie above the that they may enter a reaction relationship only at Tb703 °C, when field of aenigmatite stability. The trend of the data makes it likely that melts are strongly peralkaline. the melts would eventually intersect the aenigmatite stability curve at Under anhydrous conditions, aenigmatite has a maximum thermal lower temperatures, consistent with its occurrence in some evolved stability of 900 °C at the FMQ buffer (Lindsley and Haggerty, 1971). rocks from the volcano (Macdonaldetal.,1994). Di Carlo et al. (2010) synthesised aenigmatite in a Pantellerian The stability of aenigmatite relative to fayalite and the FeTi-oxides pantellerite at temperatures b740 °C under water-poor conditions and as a function of T and aSiO2 (Qtz) can be expressed as: below 680 °C at water-saturation. There is also some uncertainty over the : pressure range in which aenigmatite is stable. Scaillet and Macdonald FeTiO3 + 2Fe2SiO4 +Na2Si2O5 + 2SiO2 =Na2Fe5TiSi6O20 ilmenite fayalite Nds quartz aenigmatite (2001) suggested that it crystallises from peralkaline magmas only at P b100 MPa, whilst Di Carlo et al. (2010) showed that it requires At equilibrium, the variation of silica activity relative to quartz with P≥100 MPa. White et al. (2005) recognised three aenigmatite-bearing T (in K) and P (in bars) is given by: assemblages in the Pantellerian pantellerites: their assemblage 3 – equilibrated at 764 756 °C at FMQ ~0.5 to FMQ ~0.2; assemblage 4 at loga = −2878:7 = T+2:3368 + 0:288ðÞ P−1 = T−2loga −loga SiO2 Fa ilm 740–700 °C at fO2 at or just below FMQ; and assemblage 5 at b700 °C at fO just above FMQ. Aenigmatite stability appears, therefore, to be − : 2 logaNds + logaaen influenced by T–P–fO2 and its appearance cannot be used to estimate any single parameter. We now try to contribute to the debate by Aenigmatite–ilmenite equilibria (White et al., 2005) as a function examining aenigmatite stability as functions of temperature and silica of T and fO2 can be recalculated in terms of silica activity: activity. White et al. (2005) described the stability of aenigmatite relative loga = −5686:7 = T−0:2155 + 0:055ðÞ P−1 = T+0:25logf O SiO2 2 to ilmenite as a function of T and fO2: − : − : : 0 50logahem 0 25logailm ðÞ : Na2Si2O5 + 4SiO2 + 2Fe2O3 + FeTiO3 =Na2Fe5TiSi6O20 +O2 The results are plotted in T-aSiO2 (Qtz) space in Fig. 14, with melt quartz ilmenite s:s: aenigmatite gas P=1500 bars and ΔFMQ −0.5 (Pantelleria conditions; White et al., 2005). The aenigmatite-phyric Pantelleria rocks lie close to or in the We employ new thermodynamic data as follows. Free energy and aenigmatite stability field, whereas the Menengai trachytes seem to molecular volume data for ilmenite, hematite, ulvöspinel, magnetite, have too low silica activity to crystallise the phase. fayalite and quartz are from Robie and Hemingway (1995) and for Nds Taking the information in Figs. 13 and 14 together, it appears that

(melt) from Chase et al. (1985). The molecular volume for aenigmatite aenigmatite stability is a function of both fO2 and aSiO2. Although the (22.49 J/bar) was calculated from cell parameter data reported by Grew stability ﬁeld of aenigmatite expands at lower fO2,italsorequires et al. (2008). Free energies for aenigmatite were calculated by the relatively high silica activities, particularly at TN750 °C, and is overall method of Chen (1975), using free energy data from Chase et al. (1985) more stable at lower temperatures. In peralkaline silicic melts crystal- and Robie and Hemingway (1995). Activities of oxide components were lisation of aenigmatite may be favoured over that of fayalite at lower silica calculated following Andersen and Lindsley (1988); for purposes of activity. discussion, activities of unity were assumed for Nds, aenigmatite and quartz. QUILF results for Menengai and Pantelleria samples (calculated 5.8. Amphibole at P=1500 bars) are plotted in Fig. 13, with aenigmatite–ilmenite and aenigmatite–magnetite curves (also calculated at P=1500 bars) for Amphibole has been recorded as a rare phenocryst phase in oxide compositions within the ranges of the natural samples (Xusp 74– Menengai rocks by Leat et al. (1984) and Macdonald et al. (1994). 78 for Menengai and Xilm 93–97 for Pantelleria). Pantelleria samples, New analyses are available only of a crystal in K1A34/1 (Supplementary

Fig. 13. Oxygen fugacity, expressed as ΔFMQ, plotted against temperature for Menengai and Pantellerian samples. The simpliﬁed assemblages in three Pantellerian subsuites are shown by different symbols (White et al., 2005). The solid and dashed curves are relevant to the compositions of the natural FeTi-oxides (Xusp 74–78 for Menengai and Xilm 93–97 for Pantelleria). Author's personal copy

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Table 2). In the Leake et al. (1997, 2004) classification scheme, it is from the Azores (Graciosa, White et al., 1979; and São Miguel, ferrorichterite with a discontinuous narrow rim of arfvedsonite. Leat Schmincke and Weibel, 1972) and comendites from Terceira, Azores (1983) presented two arfvedsonite analyses in a pre-caldera lava. (Schmincke and Weibel, 1972) and the Olkaria complex, Kenya With the exception of the Menengai trachytes, amphibole phenocrysts (Macdonald et al., 1987; Marshall et al., 2009). None has been found in have been recorded only from peralkaline rhyolites with PIN1.25 (Fig. 12). the Menengai rocks. In the Azores rocks, the mica coexists with Its composition is variable. Noble (1965) reported Fe-rich barkevikite in a clinopyroxene+FeTi-oxide±olivine, whilst at Olkaria it most com- pantelleritic ash flow tuff from southern Nevada. Ferrorichterite was monly coexists with amphibole±olivine±FeTi-oxide. The reasons recorded in pantellerites from Pantelleria (Di Carlo et al., 2010; Nicholls for the restricted stability field are unknown. Scaillet and Macdonald and Carmichael, 1969; White et al., 2005)andEburru(Ren et al., 2006). (2001) synthesised biotite in an Olkaria comendite, showing that at

Scaillet and Macdonald (2006b) synthesised arfvedsonite at 150 MPa 150 MPa and fO2 ~FMQ, its crystallisation pointed to temperatures as under water-rich conditions and fO2=NNO−2 from an Eburru pantel- low as 660 °C under near water-saturated conditions (melt water lerite. Marshall et al. (2009) found that the composition of amphibole content approaching 6 wt.%). Such conditions are perhaps rarely met phenocrysts in Olkaria comendites broadly changes with whole-rock in peralkaline systems. However, it is unlikely that the comenditic composition, from kataphorite to richterite to silica-poor riebeckite. In trachytes from the Azores equilibrated at such low temperatures. some samples, there is variation in core composition, one example including potassian ferrorichterite, ferrikataphorite and ferrobaroisite. In 5.10. Apatite another sample, the range is magnesiorichterite, ferrorichterite and magnesio-arfvedsonite. The Pantellerian pantellerite studied experimen- Representative chemical analyses are given in Table 6 and the full tally by Di Carlo et al. (2010) contains ferrorichteritic phenocrysts; the data set may be found in Supplementary Table 1. The dominant amphibole synthesised was Fe–eckermannite–arfvedsonite, the Na components of the M site are Ca and rare earth elements (REE). The content increasing with increasing melt PI. total content of REE2O3 (including Y) varies from 1.4 to 9.2 wt.%, the Thus, in contrast to other mafic phases, especially fayalite, aenigma- latter value representing 6 mol% filling of the site. The dominant REE tite and FeTi-oxides, amphibole is compositionally diverse in peralkaline is Ce. The Fe contents (up to 0.24 a.p.f.u.) of the Menengai apatites are rhyolites. Several factors can affect its stability and composition. It has relatively large, which may have been facilitated by low fO2 conditions been shown experimentally that the Ca content of amphibole increases of crystallisation. The only significant component, other than P, in the strongly with melt water content, independent of pressure (Di Carlo Z site is Si, which forms up to 0.55 a.p.f.u., ~9% of the site. The REE are et al., 2010; Scaillet and Macdonald, 2003) and that the Na content positively correlated with Si and there is solid-solution substitution increases with increasing PI (Di Carlo et al., 2010). Replacement of towards britholite-(Ce) to a maximum of ~7 mol%. Reverse zoning in a structural water by F will increase the thermal stability of amphibole crystal in K1A34/1, with a core to rim decrease in the britholite (Conrad, 1984). Overall, however, the presence of ferrorichterite in component from 1.2 to 0.45, may possibly be ascribed to magma Menengai sample K1A34/1 is consistent with magma crystallisation at mixing, as recorded in apatite in the Olkaria rhyolites by Macdonald et low temperatures (b700 °C?), pressures N150 MPa, under water-rich al. (2008). The sum of anions in the X site ranges from 1.8 to 2.6 conditions and at fO2=FMQ(Di Carlo et al., 2010; White et al., 2005). (theoretical sum 2.0). The site is dominated by F and the phases are There is no clear relationship between the occurrence of amphibole clearly fluorapatite. and aenigmatite in the Menengai suite. They are both rare phenocryst Reports of phenocrystic apatite in peralkaline trachytes and rhyolites phases and have been found to coexist only in one pre-caldera lava are rather uncommon, possibly a result of its being overlooked due to its (Leat, 1983). Of the two, the presence of aenigmatite is better correlated small size. Mahood and Stimac (1990) and Ren et al. (2006) found it in with strongly peralkaline melts at the volcano, which is consistent with pantellerites from Pantelleria and Eburru, respectively. Schmincke and the wider data set (Fig. 12). Weibel (1972) reported it from Azores lavas and Macdonald et al. (2008) recorded it in Olkaria comendites. Apatite apparently can 5.9. Biotite crystallise over the complete spectrum of whole-rock PI values. Studies of the composition of apatite in such systems are rare, mainly because

Biotite is a rare phenocryst phase in peralkaline silica-over- P abundances are very low (P2O5 normally b0.1 wt.%; Macdonald and saturated rocks, having been recorded only in comenditic trachytes Bailey, 1973), precluding signiﬁcant apatite crystallisation. Mahood and

Fig. 14. Silica activity plotted against temperature for Menengai and Pantellerian samples. Simpliﬁed assemblages in Pantellerian subsuites (White et al., 2005) are shown by different symbols. Pressure was taken to be 1500 bars and Xilm to be 0.95. Author's personal copy

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Stimac (1990) published analyses in peralkaline trachytes and pantel- Temperatures were determined from clinopyroxene–olivine equilibria lerites from Pantelleria and Macdonald et al. (2008) reported on the from the following reactions: composition of apatite from the benmoreite–comendite series of the Olkaria complex. The compositional variation in the Menengai apatite is FoðolÞþFsðcpxÞ¼FaðolÞþEnðcpxÞ more restricted than at Olkaria; maximum REE+Si values are 2.3 and 4.0, respectively. The Pantelleria rocks show even lower enrichment in FaðolÞþHdðcpxÞ¼KstðolÞþFsðcpxÞ the britholite component, with REE+Si≤0.5 a.p.f.u., but this may partly be a result of the analyses not including HREE data. With reasonably FoðolÞþDiðcpxÞ¼MtcðolÞþEnðcpxÞ complete data available only for the Olkaria and Menengai suites, it is not yet possible to relate the variations in apatite composition to any where Fo, Fa, Kst and Mtc are the forsterite, fayalite, kirschsteinite and specific factor. A fruitful line of enquiry might, however, be the varying monticellite components in olivine, and Fs, En, Hd and Di are the F/Cl ratios in the host rocks, britholite enrichment being stronger in ferrosilite, enstatite, hedenbergite and diopside components in clino- magmas with FNCl (Olkaria and Menengai) than in those with ClNF pyroxene. Temperatures determined from these equilibria are pressure- (Pantelleria). dependent, as is aSiO2 (Qtz), which, with temperature, can be calculated from the following reactions: 5.11. Sulphides EnðcpxÞ¼FoðolÞþSiO The sulphide phase in RJK6 and M8 is pyrrhotite, with a 2 composition of 100 mol% FeS. Crisp and Spera (1987) recorded scarce ð Þ¼ ð Þþ : Fs cpx Fa ol SiO2 pyrrhotite microphenocrysts in rhyolites from the Tejeda volcano, Gran Canaria, and Mungall and Martin (1996) noted rare pyrite Oxygen fugacity can in turn be determined from equilibria phenocrysts in pantellerites of the Pico Alto volcano on Terceira Island, between iron silicates and spinel from the following reactions: Azores. Lowenstern et al. (1993) reported pyrrhotite and molybdenite in pantelleritic trachytes and pantellerites from Pantelleria; White et al. ð Þþ ¼ ð Þþ ð Þ 6Fa ol O2 2Mgt spl 3Fs cpx (2005) found pyrrhotite in the same rocks. Ren et al. (2006) recorded, from the Eburru complex, pyrrhotite±pyrite in pantelleritic trachytes ð Þþ ¼ ð Þþ : 3Fa ol O2 2Mgt spl 3SiO2 and pyrrhotite in pantellerites. On the other hand, sulphide phenocrysts were not reported from other pantelleritic suites, such as those of Mayor where Mgt is the magnetite component in spinel. All results presented Island, New Zealand (Ewartetal.,1968), Socorro Island, Mexico (Bohrson in Table 7 were calculated for a pressure of 1500 bars. and Reid, 1997) and the Gedemsa volcano, Ethiopia (Peccerillo et al., With the exception of K1A13, the results are fairly consistent. We 2003). It is not clear, however, whether the absence in these cases is real noted earlier that K1A13 is rather anomalous compositionally, which or due to the lack of a systematic search. we ascribed to post-emplacement modifications. Given the modelling The apparent scarcity of modal sulphide at Menengai and in other results, however, it may be that this rock and the other pre-caldera peralkaline suites may be due to two effects: first, sulphur is highly lavas form a quite different lineage (or lineages) to the younger soluble in peralkaline melts compared to their metaluminous extrusives. In the other samples, temperature estimates are in the counterparts and in most suites sulphide saturation has not been range 870–854 °C, which is rather lower than the range (991–888 °C) reached, and second, sulphur has been lost in a volatile phase at an established by White et al. (2005) for peralkaline rhyolites of similar earlier stage of magmatic evolution (Scaillet and Macdonald, 2006a). PI (b1.31) from Pantelleria but higher than those for pantelleritic trachytes from Eburru (793–709 °C; Ren et al., 2006). The differences 6. Discussion may be related to different melt water contents; Pantellerian magmas

have ~4–4.5 wt.% H2O(Di Carlo et al., 2010; Gioncada and Landi, 6.1. Calculation of intensive parameters 2010; Kovalenko et al., 1988), whereas in Eburru pantellerites melt water contents may be as high as 6 wt.% (Scaillet and Macdonald,

Temperature, oxygen fugacity (fO2) and silica activity (relative to 2006a, 2006b). Log fO2 values are −14.92 to −14.97, i.e. ΔFMQ −1.6 quartz, aSiO2 (Qtz)) were calculated for the Menengai trachytes using to −1.7. These values are generally lower than those for Pantellerian QUILF95, a programme that evaluates mineral equilibria in the Ca-QUIlF (FMQ −1 to FMQ −0; White et al., 2005) and Eburru (FMQ +0.5 to system (Andersen et al., 1993). The results are given in Table 7. FMQ −1.6; Ren et al., 2006) rocks. Values of aSiO2, 0.60–0.66, are

Table 7 Results of QUILF calculations, Menengai trachytes.

Pressure=1500 bars

Sample Spinel Olivine Augite Results

N–Ti N–Mg N–Mn Fo La En Wo T (°C) logfO2 ΔFMQ aSiO2 (Qtz) H2 Input 0.755 0.037 0.057 0.166 0.018 0.232 0.449 Calc 0.034 0.209 855 −14.92 −1.59 0.66 K1A13 Input 0.735 0.009 0.048 0.032 0.010 0.093 0.443 Calc 0.010 0.056 796 −15.96 −1.36 0.86 M8 Input 0.779 0.056 0.053 0.221 0.010 0.259 0.448 Calc 0.047 0.253 854 −14.97 −1.61 0.64 RJK6 Input 0.783 0.034 0.060 0.201 0.012 0.230 0.448 Calc 0.043 0.236 870 −14.77 −1.72 0.64 K1A34/1 Input 0.214 0.008 0.248 0.454 (lower−Fe) Calc 0.256 811 0.62 K1A34/1 Input 0.214 0.008 0.200 0.457 (higher-Fe) Calc 0.258 803 0.60

Italics signify values in the “Input” column that were set as trial values; the values calculated by QUILF95 (Andersen et al., 1993) are in normal font in the adjacent “Calc” column.

ΔFMQ=log (fO2)−FMQ (T). In K1A34/1, lower-Fe and higher-Fe refer to the two pyroxene populations (see text). Author's personal copy

566 R. Macdonald et al. / Lithos 125 (2011) 553–568 signiﬁcantly lower than those estimated for Pantelleria and Eburru from the host rocks in the laboratory (Di Carlo et al., 2010; Scaillet and (near unity; White et al., 2005, 2009; Ren et al., 2006). Macdonald, 2003, 2006b). Amphibole stability appears to be sensitive to

White et al. (2005) showed that ﬁve subsuites within the melt composition (e.g. the Ca/Na ratio), pressure, pH2O, pF2 and fO2,such Pantellerian eruptive sequences, distinguished by different mineral that the composition of the phase which will crystallise in any given suite assemblages, evolved along different T–fO2 paths. Our new data show might be difﬁcult to predict. We concluded here, however, that the that magmas from different magmatic complexes may also follow occurrence of ferrorichterite in a Menengai trachyte pointed to magma different T–fO2 paths. crystallisation at low temperatures (b700 °C?), pressures N150 MPa, under water-rich conditions and at fO2 =FMQ. Apart from the 6.2. Mineral stability: the state of play demonstration that biotite crystallisation in Olkaria comendites indicated temperatures as low as 660 °C at a melt water content approaching 6 wt.

Many aspects of the petrology of peralkaline trachytes and %at150MPaandfO2~FMQ (Scaillet and Macdonald, 2001), little is rhyolites are markedly similar between suites. (1) It has long been known about biotite stability in peralkaline trachytes and rhyolites. known (Lacroix, 1927; Macdonald, 1974; Macdonald and Bailey, Ilmenite and titanomagnetite have been recorded over the whole 1973; Noble, 1968) that the major element variation in peralkaline spectrum of compositions, usually occurring singly (e.g. titanomag- rhyolites is very systematic. (2) Where determined, all such suites netite at Menengai) but, rarely (e.g. in Pantelleria) together. We have appear to have evolved under distinctly reducing conditions not identiﬁed the factors which stabilise one phase relative to the

(fO2 ≤ FMQ; Mahood, 1981; Métrich et al., 2006; Nicholls and other. The “no-oxide ﬁeld” of Nicholls and Carmichael (1969), where Carmichael, 1969; Ren et al., 2006; Scaillet and Macdonald, 2001, FeTi-oxides are replaced in the crystallising assemblage by aenigma- 2003, 2006b; White et al., 2005, 2009; this paper). (3) The same tite, does not exist in all suites, e.g. the large T–X range over which studies have pointed to the low temperatures to which natural aenigmatite coexists with ilmenite in Pantellerian pantellerites peralkaline rhyolites can evolve (b700 °C). (4) The great majority of (White et al., 2005). Finally, we call for systematic studies of apatite petrogenetic studies have stressed the dominant role of alkali feldspar in peralkaline silicic suites. The extent of solid substitution between in fractionating assemblages (e.g. Avanzinelli et al., 2004; Barberi et apatite and britholite has been demonstrated only at Menengai (this al., 1975; Gibson, 1972; Macdonald et al., 1987; Marshall et al., 2009; study) and at Olkaria (Macdonald et al., 2008); the results have raised Peccerillo et al., 2003; Roux and Varet, 1975; White et al., 2009). the question as to what controls REE partitioning into apatite, one Despite these similarities, magmatic suites from different volcanic possibility being the varying F/Cl ratios in the host rocks. complexes, or within the same complex (e.g. Pantelleria; White et al., The composition of the apatite will have important geochemical 2009), have evolved along different fractionation paths, as exempli- consequences, in, for example, determining the REE abundances and

fied in Fig. 1, and possibly under different P–T–Pvol conditions. On the LREE/HREE ratios in residual melts. basis of their experimental data on a Pantellerian pantellerite, Di Carlo Supplementary materials related to this article can be found online et al. (2010) cautioned against the uncritical use of their phase at doi:10.1016/j.lithos.2011.03.011. equilibrium data to infer the magmatic conditions of other peralkaline suites. We can be optimistic that calculation of these parameters for a Acknowledgements larger range of peralkaline suites will provide major insights into the factors controlling mineral stability. We gratefully acknowledge financial support through University The different evolutionary paths must reflect either variable of Warsaw grant number BSt 1536/4 IGMiP 2010. We thank an fractionating assemblages and/or different proportions of the frac- anonymous journal reviewer for helpful comments on the original tionating phases. In this study, we have attempted to explore some of manuscript. the issues which will need to be addressed before a satisfactory resolution of these alternatives can be found. Perhaps most important References amongst these issues is the nature of alkali feldspar–melt relation- – – – ships in peralkaline rocks and, in particular, the nature of any low- Andersen, D.J., Lindsley, D.H., 1988. Internally consistent solution models for Fe Mg Mn Ti oxides. American Mineralogist 73, 714–726. temperature zone in the alkali feldspar primary phase region, which Andersen, D.J., Lindsley, D.H., Davidson, P.M., 1993. QUILF: a PASCAL program to assess has been seen as a fundamental control of melt trends in peralkaline equilibria among Fe–Mg–Mn–Ti oxides, pyroxenes, olivine, and quartz. Computers – rocks. The Menengai data show that no such zone applies to all suites, and Geosciences 19, 1333 1350. Avanzinelli, R.D., Bindi, L., Menchetti, S., Conticello, S., 2004. Crystallisation and genesis at least not in any simple form. It appears, rather, that crystallisation of peralkaline magmas from Pantelleria Volcano, Italy: an integrated petrological of feldspar with compositions at or near the minimum between the and crystal-chemical study. Lithos 73, 41–69. alkali feldspar solid solution loops drives residual liquids towards Ayalew, D., Barbey, P., Marty, B., Reisberg, L., Yirgu, G., Pik, R., 2002. Source, genesis, and timing of giant ignimbrite deposits associated with Ethiopian continental flood either Na- or K-enrichment depending on the Na/K ratio of the least basalts. Geochimica et Cosmochimica Acta 66, 1429–1448. evolved salic members of the suite. Bailey, D.K., 1974. Experimental petrology relating to oversaturated peralkaline It is also important to ascertain why quartz phenocrysts crystallise volcanics: a review. Bulletin Volcanologique 38, 637–652. Bailey, D.K., Macdonald, R., 1969. Alkali-feldspar fractionation trends and the derivation early in some suites and late in others; the timing of its appearance of peralkaline liquids. American Journal of Science 267, 242–248. will influence the fractionation trends of residual melts, during Bailey, D.K., Schairer, J.F., 1964. Feldspar–liquid equilibria in peralkaline liquids — the fractional crystallisation, towards strong or more muted silica- orthoclase effect. American Journal of Science 262, 1198–1206. Bailey, D.K., Cooper, J.P., Knight, J.L., 1974. Anhydrous melting and crystallization of enrichment. We have suggested here that it is a result of the effect peralkaline obsidians. Bulletin Volcanologique 38, 653–665. – of F or the F/Cl ratio on the position of the quartz feldspar cotectic but Baker, D.R., Vaillancourt, J., 1995. The low viscosities of F + H2O-bearing granitic melts other factors, such as total pressure and pH2O, may also be influential. and implications for melt extraction and transport. Earth and Planetary Science – Thermodynamic modelling has been used to show that the stability of Letters 132, 199 211. Barberi, F., Ferrara, G., Santacroce, R., Treuil, M., Varet, J., 1975. A transitional basalt– aenigmatite in peralkaline silicic melts is at least partly dependent on fO2 pantellerite sequence of fractional crystallization, the Boina centre (Afar rift, – and aSiO2 (Qtz), expanding at lower fO2 where high silica activities are Ethiopia). Journal of Petrology 16, 22 56. relatively high, particularly at TN750 °C. Whereas the compositions of Black, S., Macdonald, R., Barreiro, B.A., Dunkley, P.N., Smith, M., 1998. Open system alkaline magmatism in northern Kenya: evidence from U-series disequilibria and olivine and clinopyroxene phenocrysts vary fairly systematically across radiogenic isotopes. Contributions to Mineralogy and Petrology 131, 364–378. the spectrum of whole-rock compositions, and those of aenigmatite and Bohrson, W.A., Reid, M., 1997. 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