Xenoliths from the sub-volcanic lithosphere of Mt Taranaki, New Zealand

Journal of Volcanology and Geothermal Research 190 (2010) 192–202

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Journal of Volcanology and Geothermal Research j o u r n a l h o m e p a g e : w w w. e l s e v i e r. c o m / l o c a t e / j v o l g e o r e s

Xenoliths from the sub-volcanic lithosphere of Mt Taranaki, New Zealand Kerstin Gruender a, Robert B. Stewart a,⁎, Stephen Foley b a b

INR, Massey University, Private Bag 11 222, Palmerston North 4442, New Zealand Institut für Geowissenschaften, Johannes Gutenberg Universtität, Becher-Weg 21, 55099 Mainz, Germany

a r t i c l e

i n f o

Article history: Received 22 August 2008 Accepted 22 September 2009 Available online 3 October 2009 Keywords: xenolith metasomatism Mt Taranaki Egmont volcano lithosphere subduction

a b s t r a c t Mount Taranaki is located 140 km west of the Taupo Volcanic Zone and represents the most westerly expression of subduction-related volcanism on the North Island of New Zealand. Taranaki is a predominantly high-K arc volcano but compositions range from basaltic andesite to andesite with minor dacite and basalt. The sub-volcanic basement under Taranaki is thought to comprise calc–alkaline plutonic and metamorphic rocks of the Median Batholith, overlain by a sequence of Cretaceous and Tertiary sediments. Taranaki lavas contain abundant xenoliths that represent samples of the upper to lower crust beneath the volcano. The xenolith suite has been initially organised into six groups based on petrography, geochemistry and inferred genetic relationships: supracrustal sedimentary rocks (1), mafic hornfels (2), garnet gneiss (3), granite and granodiorite (4), finely banded amphibolitic gneiss (5) and gabbros and ultramafic rocks (6). Groups 1, 3 and 4 are derived from the Median Batholith basement and Cretaceous–Tertiary sediments of the Taranaki Basin while Groups 2, 5 and some fine grained gabbros from Group 6 could either be derived from the Median Batholith or be cognate xenoliths. Group 6 gabbros and ultramafic rocks are dominated by clinopyroxene, amphibole and plagioclase and are predominantly cumulate in origin. The Egmont xenoliths can also be classified into the Type I and Type II xenoliths defined by Frey and Prinz (1978). Type I dunite and wehrlite xenoliths are only present in basaltic andesite host rocks and are sourced from depleted upper mantle whereas Type II xenoliths predominate in the more siliceous andesites and are sourced from the lower crust. The separate source depths for the two rock types can be explained by the “hot zone” model where the andesites have much greater interaction with the lower crust than the basaltic andesites. Some xenoliths contain glass of rhyolitic to trachyitic compositions with up to 6% K2O that represent partial melts of the sub-volcanic lower crust and may give rise to the andesite magma compositions by mixing with lower crustal residual crystals. The widespread occurrence of amphibole in the Egmont xenolith suite reflects the fluid-rich environment of arc magma systems. © 2009 Elsevier B.V. All rights reserved.

1. Introduction Xenoliths in volcanic rocks provide a window into the composition and distribution of sub-volcanic lithologies (Graham, 1987; Graham et al., 1988; O'Reilly et al., 1989; Graham et al., 1990) and sample the vertical extent of the magma plumbing system from mantle to supracrustal rocks (Wysoczanski et al., 1995). Here we focus on the deeper sourced xenoliths that provide information on lower crustal and mantle compositions beneath Egmont volcano, Taranaki, New Zealand. Much data on sub-volcanic compositions are derived from alkaline rocks which appear to rise rapidly from their mantle source and frequently sample the strata through which they pass (Griffin and O'Reilly, 1987; Wysoczanski et al., 1995; Chen and Arculus, 1995; Alletti et al., 2005). Frey and Prinz (1978) identified Type I and Type II ⁎ Corresponding author. E-mail addresses: [email protected] (K. Gruender), [email protected] (R.B. Stewart), [email protected] (S. Foley). 0377-0273/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2009.09.014

xenoliths; Type I comprise olivine-bearing ultramafic compositions with Cr-diopside as the clinopyroxene and also contain Cr-spinels. They have low Al2O3, high Mg/Fe ratios and represent depleted upper mantle, which has undergone multiple melt extractions (Kovács et al., 2004; Dessai et al., 2004). Texturally the Type I lithologies show evidence of deformation and recrystallisation in the presence of deformation lamellae, granoblastic textures and foliations (Griffin and O'Reilly, 1987; Alletti et al., 2005; Ghent et al., 2008). In contrast, Type II are uppermost mantle/lower crustal rocks (Alaugite series of Wilshire and Shervais (1975)), which have cumulate or meta-igneous textures and are cumulate residues from partial crystallisation of basaltic magmas at or near the base of the crust. Type II lithologies generally do not exhibit strain or granoblastic textures. Both Type I and II xenoliths may show evidence of metasomatism (Kovács et al., 2004; Dessai et al., 2004). A further group of deep crustal xenoliths is mafic and felsic granulites and charnockites that are inferred to comprise much of the lower crust (Kempton et al., 1990; Rudnick and Fountain, 1995; Kovács and Szabó, 2005; Ghent et al., 2008). These have foliated to

K. Gruender et al. / Journal of Volcanology and Geothermal Research 190 (2010) 192–202

granoblastic textures and show evidence of reaction between Caplagioclase and olivine to form spinel and clinopyroxene that characterises granulite facies metamorphism (Alletti et al., 2005). Xenoliths in the alkaline basalts therefore record the passage of magma through depleted upper mantle, lower crustal cumulates, granulites and on through supracrustal lithologies. The record of xenoliths from subduction zone magmas is much less comprehensive and most examples of xenoliths in subduction volcanics come from the western Pacific (Arai et al., 2007). They have been used in arcs to provide clues about magma generation, preeruptive processes and interaction with the crust, for example the Aleutian Arc (Conrad and Kay, 1984; Debari et al., 1987), the Lesser Antilles (Arculus and Wills, 1980) and Mt Ruapehu in the Taupo Volcanic Zone of New Zealand (Graham, 1987; Graham and Hackett, 1987; Graham et al., 1988, 1990), as well as characterising the composition of the mantle wedge (Arai et al., 2007). Xenoliths in arc rocks are predominantly interpreted as cognate and supracrustal in origin (Arculus and Wills, 1980; Graham et al., 1990). Because the lower crust/upper mantle zone that interacts with arc magmas is generally hotter and contains a greater melt fraction than that encountered by alkaline basalts, the lower crustal lithologies are more easily disaggregated and recorded predominantly as either single crystals or glomerocrysts (Arculus and Wills, 1980; Stewart et al., 1996; Price et al., 1999, 2005; Annen et al., 2006; Price et al., 2008). In this paper we report on a suite of xenoliths from Egmont volcano, Mt Taranaki, New Zealand which are both cognate and of lower crust/upper mantle origin. This study presents new petrographical data and major element analyses of minerals within the xenoliths and distinguishes possible origins and genetic relationships. The data also provide further constraints on subsurface geology of the Taranaki region and give unique insights into the sub-volcanic lithosphere and lower crust beneath the volcano. 2. Regional setting Egmont volcano (Mount Taranaki) is centred on the Taranaki Peninsula in the western North Island of New Zealand and about 140 km west of the Taupo Volcanic Zone (TVZ), the main region of active volcanism in New Zealand (Fig. 1). Egmont is the youngest of four volcanic centres in the region, forming a northwest–southeast trending lineament with volcanism becoming progressively younger towards the southeast. The oldest volcanic centre is Paritutu, including the Sugarloaf Islands near New Plymouth (1.75 Ma), followed by Kaitake (0.57 Ma), Pouakai (0.25 Ma) and Egmont (b0.12 Ma) (Neall, 1979; Neall et al., 1986). Volcanic rocks of Egmont are predominantly high-K andesites and basaltic andesites, with minor dacites and high-alumina basalts (Stewart et al., 1996). The youngest eruptives at the summit have the highest SiO2-content and the lavas have also become progressively more K-rich with time (Neall et al., 1986; Price et al., 1992; Stewart et al., 1996; Price et al., 1999). The volcanoes in the Taranaki region are located about 400 km west of the trench, are c. 250 km above a Wadati–Benioff zone (Adams and Ware, 1977; Boddington et al., 2004; Reyners et al., 2006) and overlie 25 to 35 km thick continental crust (Stern and Davey, 1987). However, the slab is only traceable to the southeastern part of the Taranaki region, c. 35 km east of Mt Taranaki (Stern et al., 2006) so is not clearly expressed beneath the volcanoes. In comparison, the volcanoes in the TVZ overlie only c. 15 km of relatively hot continental crust and the Wadati–Benioff zone is at 80 km depth (Stern and Davey, 1987; Stratford and Stern, 2006). Recent seismic studies of the Taranaki region show a brittle–ductile transition zone at 10 km depth beneath Mt Taranaki and 3D seismic velocity tomographic imaging of the Taranaki volcanoes shows a volcanic root system of around 5 km in diameter, extending to a depth of c. 10 km (Sherburn and White, 2005, Sherburn et al., 2006).

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Fig. 1. Partial Palaeozoic–Mesozoic map of New Zealand showing the distribution of terranes. The Brook Street, Maitai, Dun Mountain ophiolite and Murihiku Terranes, together with part of the Median Batholith (which extends through to the Cretaceous), are fossil Carboniferous to Jurassic arcs. Egmont lies near the western boundary of Median Batholith basement.

The Taranaki Peninsula is the onshore component of the Taranaki Basin and the upper 6 km of the crust comprises Cretaceous to Tertiary sedimentary rocks (King and Thrasher, 1996). The deeper basement geology of the Taranaki region has been extrapolated from the South Island and is considered to be part of the Median Batholith (Mortimer et al., 1999). This is based on information from oil exploration drill holes, seismic studies, magnetic anomalies and some information from xenoliths in volcanic rocks (Wodzicki, 1974; Knox, 1982; Gamble et al., 1994; King and Thrasher, 1996; Mortimer et al., 1997; Sutherland, 1999). The Median Batholith is made up of diorites, gabbros and granitoid rocks (Challis et al., 1994; Rattenbury et al., 1998) and some igneous xenoliths have been reported to show similarities to rocks from the Median Batholith. Large sandstone xenoliths in the Pungarehu Formation in the western Taranaki region have been correlated to the Eocene–Oligocene Kapuni Group sandstone at approximately 3.5 km depth beneath the volcano (Collen et al., 1985). However, no exploration wells close to Mt Taranaki have penetrated basement rocks and proposed basement terrane boundaries on the Taranaki Peninsula remain speculative (King and Thrasher, 1996; Mortimer et al., 1997). 3. Materials and methods Rock samples with xenoliths were collected from a variety of sites along beaches and on the mountain. Xenolith host rocks range from basaltic andesite to dacite and are typically porphyritic with crystal contents between 30 and 60% (Neall et al., 1986; Price et al., 1992; Stewart et al., 1996). All of the materials are reworked from primary lava flows and transported as debris flows and avalanches, from where it is again reworked by fluvial or beach processes. There is therefore no stratigraphic control on the samples, except that where found in debris avalanche deposits a minimum age can be ascribed. The focus in this paper is on the spectrum of compositions represented by the xenoliths as representative of sub-volcano lithologies.

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Table 1 Representative olivine and spinel analyses, Egmont xenoliths. Rock type numbers are: 1 = thermally altered sediment, 2 = amphibolitic gneiss, 3 = mafic hornfels, 4 = fine grained gabbro, 5 = granodiorite, 6 = hornblende gabbro, 7 = hornblende pyroxene gabbro, 8 = meta-hornblende pyroxenite, 9 = hornblende pyroxenite/pyroxene hornblendite, 10 = hornblendite, 11 = olivine hornblende clinopyroxenite, 12 = troctolite/clinopyroxenite, 13 = clinopyroxenite, 14 = olivine clinopyroxenite, 15 = wehrlite, and 16 = dunite. Ol = olivine, Cr-sp = chrome spinel, and Al–Mg-sp = spinel. FeO* is total Fe as FeO. Rock type

(11)

(13)

(12)

(7)

Mineral

ol

ol

ol

SiO2 TiO2 Al2O3 Cr2O3 FeO⁎ MnO MgO CaO Na2O K2O ZnO NiO V2O3 Total

37.36 nd 0.01 0.02 23.25 0.33 38.79 0.10 nd nd 0.07 0.09 nd 100.02

38.82 nd nd 0.04 20.21 2.26 39.59 0.16 0.01 nd 0.11 nd nd 101.20

39.40 nd nd 0.02 17.47 0.27 43.22 0.06 nd nd 0 0.14 nd 100.58

(15)

(16)

(15)

(12)

(olivine in ol T952x3)

ol

Cr-sp

Al–Mg-sp

39.72 nd nd 0.03 16.39 0.64 43.83 0.03 0.01 nd 0.01 0.10 0.01 100.78

40.78 0.01 0.01 0.03 9.24 0.18 49.36 0.05 nd nd 0.02 0.35 0.01 100.04

0.04 1.36 16.79 34.37 35.25 0.23 10.56 0.11 nd 0.01 0.13 0.20 0.21 99.27

0.01 0.09 62.83 0.06 19.61 0.97 16.19 nd 0.03 0.01 0.51 0.01 0.01 100.35

40.71 nd 0.02 0.02 12.02 0.16 47.41 0.04 nd nd 0.05 0.24 nd 100.67

Polished sections of the xenoliths were prepared and analysed microscopically. Mineral major element analyses were carried out by electron microprobe (Jeol Superprobe JXA 8900) at Johannes Gutenberg University (Mainz, Germany) under the following analytical conditions; acceleration voltage = 20 kV for olivine, pyroxenes, spinel, Fe–Ti oxides and 15 kV for feldspars, glass, other minerals; beam current = 8 nA for feldspars, 12 nA for all other minerals; beam diameter = 5 µm for feldspars and glasses, 2 µm for all other minerals. The xenolith rock types were classified according to the IUGS recommendations for igneous rocks (Le Maitre, 1989). Amphiboles and pyroxenes were classified using the IMA nomenclature (Leake, 1997 and Morimoto, 1988) and amphibole analyses were recalculated using the procedure outlined by Schumacher (1997). 4. Results

also ubiquitous in thin section and comprise similar mineral assemblages to those in the macroscopic xenoliths. The contact relationships between xenoliths and host rock vary, and no systematic textures have been identified for specific xenolith rock types. Some have sharp angular contacts to the host rocks, while others have a reaction rim of amphibole and/or titanomagnetite and clinopyroxene or show glass-filled fractures and melt areas within them. Xenoliths are grouped primarily according to their mineralogy and textures, taking into account some mineral chemistry and inferred genetic relationships. Most xenolith types can be distinguished and grouped macroscopically; only very fine grained xenoliths require microscopic analysis because these can appear very similar in hand specimen (e.g. mafic hornfels and meta-sediments). Estimated pressures, giving depth constraints, obtained using Al in hornblende geobarometry (Hollister et al., 1987) were taken into account in the final grouping for amphibolitic gneiss, hornblende pyroxene gabbro and ultramafic rock types. To keep the list of xenolith groups short, some simplification was applied to the variety of rock textures found, for example the degree of foliation and grain size. Representative analyses for xenolith minerals are listed in Table 1 (olivine and spinel), Table 2 (amphibole and biotite) and Table 3 (clinopyroxene). 4.1.1. Supracrustal sedimentary rocks Sedimentary type xenoliths typically occur as small (10–50 mm diameter), angular clasts. Lithologies range from mature quartz-rich (95%) sandstones to less mature sediments containing up to 35% subrounded feldspars and clinopyroxene fragments in a clay-rich matrix. Coarse and fine grained rock types occur, with very fine grained varieties often being partially melted to form silica-rich (c. 80% SiO2) glass (Table 4) and Ca-rich clinopyroxenes. Similar glasses have been reported from sedimentary xenoliths in Ngauruhoe lavas (Graham et al., 1988). Creamy-white coloured and relatively unaltered, friable quartz sandstone xenoliths from the Kapuni Group occur in the Pungarehu Formation debris avalanche. Kapuni Group quartz sandstone does not outcrop in the Taranaki region and oil well data show it to occur at 3.6 km depth beneath the volcano (Collen et al., 1985). 4.1.2. Garnet gneiss These xenoliths are very rare and show a typical gneissose texture. The mafic layers contain almandine garnet, together with amphibole and biotite. Garnets are zoned with Mg-rich cores and more Fe-rich rim compositions. The felsic layers are rich in quartz and feldspar.

4.1. Petrography and mineral chemistry of xenoliths Xenoliths range in lithology and size from 10 mm up to 500 mm in diameter and are common in all Taranaki lavas. Glomerocrysts are

4.1.3. Granite and granodiorite These rock types are also rare as xenoliths and typically comprise varying amounts of plagioclase (An52–56), alkali feldspar and quartz,

Table 2 Representative amphibole (amph) and biotite (bio) analyses, Egmont xenoliths. FeO* is total Fe as FeO. O = F, Cl is an adjustment for excess oxygen calculated due to the presence of F and Cl. Rock type numbers as for Table 1. Rock type

(8)

(5)

(7)

(2)

(6)

(11)

(9)

(10)

(9)

Mineral

amph

amph

amph

amph

amph

amph

amph

amph

amph

bio

SiO2 TiO2 Al2O3 Cr2O3 FeO⁎ MnO MgO CaO Na2O K2O F Cl Total O = F,Cl Total − F,Cl

52.85 0.48 4.03 0.26 10.37 0.29 18.08 12.44 0.97 0.26 nd nd 100.03

48.32 1.05 6.64 0.03 13.93 0.54 14.36 11.97 1.18 0.47 0.02 0.06 98.57 0.02 98.55

47.57 1.44 7.13 0.08 12.78 0.28 14.75 12.01 1.49 0.74 0.11 0.05 98.43 0.06 98.37

46.82 1.66 7.61 0.17 11.76 0.18 15.58 11.69 0.83 1.70 0.24 0.01 98.25 0.10 98.15

45.84 1.30 9.32 0.05 12.79 0.44 14.83 11.22 2.15 0.41 nd 0.05 98.39 0.01 98.38

41.62 2.12 12.00 0.15 12.62 0.21 14.32 11.89 2.42 0.80 nd 0.05 98.20 0.01 98.19

40.25 2.61 13.35 0.06 12.29 0.22 13.19 12.03 2.37 0.79 0.03 0.02 97.20 0.02 97.19

41.59 1.92 13.91 0.11 12.37 0.17 13.11 11.87 2.58 0.70 nd 0.01 98.33

40.96 1.97 14.24 0.04 10.86 0.10 14.08 12.18 2.60 0.97 nd nd 98.00

98.33

98.00

37.10 4.32 14.59 0.06 12.35 0.16 16.92 0.03 1.10 8.39 0.25 0.05 95.33 0.12 95.21

100.03

(2)

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Table 3 Representative clinopyroxene analyses, Egmont xenoliths. Mg# = magnesium number, En = enstatite. Rock type numbers as for Table 1. FeO* is total Fe as FeO. Rock type

(1)

(3)

(4)

(2)

(7)

(9)

(11)

(12)

(14)

(15)

Mineral

cpx

cpx

cpx

cpx

cpx

cpx

cpx

cpx

cpx

cpx

SiO2 TiO2 Al2O3 Cr2O3 FeO⁎ MnO MgO CaO Na2O K2O NiO Total Mg# En

51.08 0.07 0.65 nd 8.67 0.55 12.61 24.88 0.13 nd 0.01 98.68 0.72 0.36

51.77 0.12 1.07 0.04 14.00 0.11 8.97 23.66 0.23 nd 0.02 100.16 0.53 0.26

51.23 0.11 1.57 0.04 12.95 0.34 9.95 23.61 0.45 nd nd 100.38 0.58 0.28

51.76 0.23 2.04 0.06 8.00 0.30 14.29 22.43 0.29 nd 0.03 99.52 0.76 0.4

52.53 0.36 1.64 nd 7.74 0.41 14.45 22.55 0.42 nd nd 100.24 0.77 0.4

48.1 0.97 6.30 0.02 7.14 0.12 13.14 23.47 0.33 nd 0.02 99.75 0.77 0.36

52.89 0.22 1.42 0.07 7.3 0.33 15.54 21.32 0.37 0.01 nd 99.53 0.79 0.43

53.03 0.13 1.35 0.02 4.54 0.54 14.89 24.88 nd nd nd 99.39 0.85 0.41

53.03 0.25 1.95 0.56 4.29 0.12 17.02 22.13 0.18 nd 0.03 99.59 0.88 0.47

54.05 0.24 1.89 0.69 3.24 0.06 16.70 23.20 0.29 nd 0.03 100.42 0.9 0.45

with green magnesiohornblende and minor amounts of brown biotite. Accessory minerals include apatite, zircon and titanite. 4.1.4. Mafic hornfels These xenoliths are light green to light brown in hand specimen (very similar to some sedimentary xenoliths) and very fine grained to microcrystalline. In thin section they show a granoblastic texture and the mineral assemblage clinopyroxene+ calcic plagioclase (An80–93) + titanite + some sulphides, Fe–Ti oxides and minor amounts of wollastonite. Mineral grains are between 25 and 40 µm in diameter, but larger (400–800 µm) relict grains of plagioclase and clinopyroxene are present as well (Fig. 2a). These are usually spongy or sieve textured and are often embayed. Relict plagioclase has the same compositional range as plagioclase in gabbros (An80–90). Clinopyroxenes are diopsides, but differ from most other xenolithic clinoproxenes observed by having a higher Fe content. Sulphides and Fe–Ti oxides occur mainly in areas close to host rock contacts. Cross cutting sulphide veins occur in some samples, as well as veins filled with plagioclase (An93). 4.1.5. Finely banded amphibolitic gneiss These xenoliths are typically very fine grained and strongly foliated with alternating felsic and mafic layers. The individual layers are between 0.5 and 1 mm thick and irregular banding, schlieren and small folds are common (Fig. 2b). Mafic layers are made up of 50–70% brown amphiboles (edenite, magnesiohastingsite and titanian magnesiohastingsite), 10–30% plagioclase (An58–75) and clinopyroxene (diopside to augite), and 5– Table 4 Representative glass analyses, Egmont xenoliths. Rock type numbers as for Table 1. FeO* is total Fe as FeO. Glass type

SiO2 TiO2 Al2O3 Cr2O3 FeO⁎ MnO MgO CaO K2O Na2O BaO Cl F O = F,Cl Total

(11)

(7)

Trachytic

Rhyolitic

(1) Rhyolitic

63.25 0.612 18.83 nd 1.89 0.166 0.39 1.122 5.54 6.39 0.235 0.207 0.159 98.791 0.114 98.677

73.94 0.267 12.79 0.071 1.77 0.097 0.039 0.902 5.32 3.4 1.52 0.086 nd 100.202 0.019 100.183

80.47 0.1 8.63 0.047 1.48 nd 0.037 0.486 4.88 2.21 0.186 0.136 nd 98.662 0.031 98.631

10% titanomagnetite. Some ilmenite and brown biotite are also present. Larger amphibole grains are often embayed and decomposed to clinopyroxene, plagioclase and titanomagnetite. The felsic layers consist of up to 50% plagioclase, 30–40% clinopyroxene, 5–10% apatite and minor amounts of titanomagnetite and amphibole. Occasionally sulphide grains occur and the relatively thick felsic layers often have clear glass in the centre of the layers and around mineral grains. Clear glass is sometimes also found in fractures that cut through the xenolith from the surrounding volcanic host. 4.1.6. Gabbros and ultramafic rocks The xenoliths included in this group are the most common types found in Mt Taranaki lavas. The term “gabbro” is here used to describe medium to coarse grained xenolith lithologies where the main mineral phases are plagioclase, amphibole and clinopyroxene. Although amphibole may be more abundant than clinopyroxene it clearly replaces clinopyroxene and is not primary. Xenoliths are described as “ultramafic” when containing less than 10% modal plagioclase. Amongst all studied xenoliths, the most abundant rock type is medium to coarse grained gabbro (~60%), followed by ultramafic rocks (~ 30%) that typically appear dark green to dark brown in hand sample. Occasionally composite xenoliths occur, for example ultramafic types adjacent to or enclosed by a gabbroic xenolith. Textures range from completely non-foliated to strongly foliated and the foliation is interpreted as a primary igneous feature rather than metamorphic because of the presence of cumulate textures and absence of granoblastic textures or strain-induced recrystallisation. 4.1.6.1. Hornblende pyroxene gabbro. Although clinopyroxene is an essential mineral in the definition of gabbro, it is used here to differentiate this xenolith type from hornblende gabbro. Hornblende pyroxene gabbro xenoliths occur as fine grained to coarse grained varieties. The modal mineral content is variable with 30–60% plagioclase (An80–90), 20–35% brown amphibole (mainly titanian magnesiohastingsite), 10–35% clinopyroxene (diopside), 3–15% titanomagnetite with accessory minerals apatite and zircon. Textures are gabbroic and a very common feature in many rocks is the replacement of clinopyroxene by amphibole. Clinopyroxene grains are mantled by brown amphibole and often embayed and shaped irregularly (Fig. 2c). Some xenoliths contain small ultramafic inclusions. The minerals are typically more coarse grained than their gabbro host and can be fragmented. One particular xenolith, T95 2 × 3, shows even more complex textures which indicate metasomatism of an original cumulate and subsequent decompression reaction of the amphibole. The xenolith contains oscillatory zoned plagioclase, as well as unzoned and normally zoned plagioclase and different sized

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Fig. 2. Egmont xenolith fabrics and reaction textures. a) Relict plagioclase grain within assemblage of very fine grained cpx + plag+ tmt in mafic hornfels (cross polarised light). b) Irregular folds and schlieren structures in amphibolitic gneiss (plane polarised light). c) Typical hornblende pyroxene gabbro, embayed cpx being replaced by brown amph (plane polarised light). d) Hornblende pyroxene gabbro T95-2x3, olivine crystal cluster with surrounding layer of px and tmt, rimmed by layer of bio and amph (plane polarised light). plag = plagioclase, ol = olivine, px = pyroxene, cpx = clinopyroxene, opx = orthopyroxene, amph = amphibole, bio = biotite, sp = spinel, tit = titanite, tmt = titanomagnetite, and ap = apatite.

amphibole and clinopyroxene grains. Most clinopyroxenes are mantled by amphibole, but some larger amphibole grains also show reaction to clinopyroxene and titanomagnetite along their edges. In the same xenolith small patches of clear or brown glass are present close to some large amphiboles and a few crystal clusters of olivine occur. The olivine is mantled by a layer of small clinopyroxenes, orthopyroxenes and titanomagnetite, which are in turn mantled by biotite and amphibole (Fig. 2d). 4.1.6.2. Hornblende gabbro. Hornblende gabbro xenoliths are less common than hornblende pyroxene gabbro xenoliths and contain 40– 50% plagioclase (An59–74), 40–60% green (magnesiohornblende to edenite) or brown (edenite) amphibole, 5–15% titanomagnetite and rare ilmenite. Accessory mineral phases are apatite, titanite and sulphides. In some xenoliths, areas with different grain sizes occur and in one of these glass films occur around grains in the transition area between fine and coarser grained zones. 4.1.6.3. Hornblendite and clinopyroxenite. The majority of these ultramafic xenoliths are hornblende pyroxenites and pyroxene hornblendites. Modal contents of either mineral range between 20 and 60% with additional 8–20% titanomagnetite. Some xenoliths

contain up to 5% plagioclase, predominantly present as small, interstitial grains. Fine grained and coarse grained cumulate rock textures occur and one particular hornblendite xenolith is made up of large amphibole crystals up to 10 cm long. In some hornblendite and clinopyroxenite xenoliths, clinopyroxene and amphiboles appear as an irregular intergrowth, in others clinopyroxene and titanomagnetite are poikilitically enclosed by amphibole (Fig. 3a). Chemically, amphiboles are titanian pargasite and pargasite; the coarse grained varieties have the highest Mg numbers observed and also show the highest Al2O3 contents. The clinopyroxene is diopside. 4.1.6.4. Meta-hornblende pyroxenite. This type includes an unusual xenolith with textures that resemble hornblende pyroxenites but is quite different in terms of mineral assemblage, chemistry and microtextures. In hand specimen the xenolith is mostly very fine grained and light green with a visible network of irregular felsic veins, some of which contain apparently brecciated larger crystals of amphibole and clinopyroxene. In thin section these have been partially broken down to form smaller amphibole and clinopyroxene crystals. The light green areas in the rock comprise medium to coarse diopsidic clinopyroxenes (Fig. 3b) that have been partially or completely replaced by brown amphibole and in parts also replaced

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Fig. 3. Metasomatic reaction textures and troctolite fragment in Egmont xenoliths. Abbreviations as in Fig. 2. a) Amph poikilitically enclosing subhedral cpx and tmt in hornblende pyroxenite (plane polarised light). b) Meta-hornblende pyroxenite, showing coarse grained cpx partially replaced by amph and partially broken down to finer grained px towards the more felsic areas (plane polarised light). c) Troctolite patch within fine grained clinopyroxenite. Plag shows even polysynthetic twinning (cross polarised light). d) Green Al-sp “spots” within clinopyroxenite. Al-sp is enclosed by poikilitic plag, ap and cpx (plane polarised light).

by small subhedral orthopyroxene and clinopyroxene. Amphiboles are mainly magnesiohornblende, but some smaller grains are edenite. The felsic veins contain plagioclase (An56) and some clinopyroxene and titanomagnetite. In some parts of the xenolith, clear glass with euhedral to subhedral small ortho- and clinopyroxene, amphibole and plagioclase crystals is present. 4.1.6.5. Olivine-bearing ultramafic xenoliths (olivine hornblende clinopyroxenite and troctolite). Olivine-bearing rocks are rarely found as xenoliths in Mt Taranaki lavas and only in association with other xenolith types such as gabbros and hornblende pyroxenites. An exception is olivinebearing xenoliths in basaltic andesite and these are described as a separate group. Olivine hornblende pyroxenite contains up to 30% modal olivine, but the olivine is not evenly distributed. Instead, medium grained and inclusion-rich clinopyroxenes in hornblende pyroxenite have been partially replaced by amphiboles and in places grade into a finer grained assemblage comprising inclusion-free clinopyroxene, olivine and titanomagnetite. One xenolith is zoned with a fine grained core containing olivine, clinopyroxene and amphibole surrounded by coarser grained clinopyroxene and amphibole towards the host rock contact. The finer grained clinopyroxenes have less CaO and slightly less Al2O3 than medium and coarse grained clinopyroxenes in same

xenolith. Olivines are the most iron rich (Fo73–75) amongst the studied rocks and contain the least amounts of NiO (b0.01 wt.%) of those analysed. Another xenolith containing olivine is troctolitic, occurring as irregular shaped but discrete coarser grained areas within very fine grained clinopyroxenite (Fig. 3c). Green Al-spinel occurs throughout the surrounding clinopyroxenite in “spots” and larger, subhedral grains are enclosed poikilitically by plagioclase. In contact with the clinopyroxenite, a greyish rim of clinopyroxene and apatite is developed (Fig. 3d). Oxide phases in troctolite are magnetite and ilmenite. Within the clinopyroxene-rich parts of the rock, ilmenite and sulphides occur. Plagioclase in troctolite and around spinel is almost pure anorthite (An98–100). Clinopyroxenes are Ca-rich diopside and have high Mg numbers (0.83–0.86), but relatively low Cr2O3 (0– 0.07 wt.%); especially compared to wehrlite and pyroxenite xenoliths from basaltic andesite lavas. Olivines are Fo77–78, very Ca-rich (N0.15 wt.%), with very low NiO concentrations. 4.1.6.6. Xenoliths in basaltic andesite. These xenolith types have been grouped separately because they solely occur in basaltic andesite host rocks. Basaltic andesite contains many angular fragments of other andesite lithologies, together with abundant small, olivine-bearing ultramafic xenoliths, xenocrysts and olivine-clinopyroxene crystal clusters.

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Type I (Frey and Prinz, 1978) pyroxenites and dunites are rare and only found in basaltic andesites. The most primitive xenoliths found in Egmont lavas are dunite and wehrlite and their mineral assemblage of olivine + Cr-diopside+ Cr-spinel confirms that they are Type I. Dunite comprises 95% olivine (Fo90), 2–4% clinopyroxene (diopside) and 1–2% small Cr-spinel inclusions in clinopyroxene. Wehrlite and olivine pyroxenite have the same mineral assemblages, but different modal abundances (up to 20% clinopyroxene in wehrlite and up to 10% olivine in olivine pyroxenite) and Cr-spinel is not always present. Olivine in wehrlite is Fo87 and Fo82 in olivine clinopyroxenite. The olivines are highly magnesian (FoN80) and also have high NiO contents (N0.25%). Olivine in dunite has the highest NiO contents (0.35–0.4 wt.%), followed by c. 0.25 wt.% NiO in wehrlite and between 0.1% and 0.2 wt.% NiO in olivine clinopyroxenite. Clinopyroxenes in dunite and wehrlite have the highest Mg numbers (N0.88) and also the highest Cr-contents (up to 0.70 wt.% Cr2O3) in the Egmont xenoliths. Cr-contents in clinopyroxenes from olivine pyroxenite and pyroxenite are lower, as are NiO and Fo-contents in olivines in the same xenoliths. Olivine and clinopyroxene react to form pargasitic amphibole in the amphibole-bearing xenoliths. Multiple generations of amphibole and clinopyroxene are present in some rocks, reflecting a complex history of metasomatism. Fine grained gabbroic xenoliths consist of up to 50% Fe-rich clinopyroxene, 45% plagioclase and 5% titanite. The mineral assemblages, textures and chemistries are nearly the same as for mafic hornfels xenoliths, but lack relict mineral grains and sulphides. Further, plagioclase in the mafic hornfels is An80–90 whereas it is An64–69 in fine grained gabbro. 4.2. Glass in xenoliths Clear or brown glass occurs in some xenolith samples (Table 4). It can be found in fractures that cross-cut xenoliths, in areas close to the contact with the volcanic host rock, between mineral grains (olivine hornblende pyroxenite, meta-hornblende pyroxenite), as light brown films around grains in the contact area between fine and coarse grained areas within one rock type (e.g. hornblende gabbro) or in felsic rock layers (finely banded amphibolitic gneiss). One thermally metamorphosed sedimentary xenolith contains up to 50% clear glass with small euhedral crystals of plagioclase and clinopyroxene. This glass has the highest SiO2 concentrations (~ 80 wt.% SiO2) of all analysed glasses. Most glasses are rhyolitic in composition and rich in K2O (4.1–6.2 wt.%); the total range of silica is from 62.9 to 80.6 wt.% SiO2, with most analyses falling between 69 and 74 wt.% SiO2. Trachytic glass occurs as inclusions in phenocrysts in basaltic andesites and interstitial glass between mineral grains in olivine hornblende pyroxenite (Fig. 4). 4.3. Pressure conditions of xenoliths The geobarometer based on total Al contents in amphibole was applied to amphibole-containing xenoliths (Hollister et al., 1987). The barometer is experimentally calibrated to the mineral assemblage quartz + K-feldspar + plagioclase + biotite + Fe-Ti oxide + titanite over a pressure range of 2–8 kbar with an estimated error of ±1 kbar. The required conditions for this geobarometer are not met for Taranaki rocks as K-feldspar was not identified microscopically in the xenoliths and quartz was only present in some; the calculated pressures may therefore be overestimated by up to 1.5 kbar (Hollister et al., 1987; Johnson and Rutherford, 1989; Anderson and Smith, 1995). The results show the range of pressure estimates for the different xenolith groups (Table 5). 5. Discussion The xenolith suite from Egmont volcano comprises a diverse set of rocks that represent the variations in lithology of the crustal sequence

Fig. 4. Composition of glasses from xenoliths showing three distinct compositional groups. High silica partial melts are from sediment xenoliths, gabbroic and gneissic xenoliths contain glasses of rhyolite composition while trachytic glasses occur in ultramafic hosts. A glass inclusion in cpx from a basaltic andesite is also trachytic.

beneath the volcano. Most are Type II xenoliths (Frey and Prinz, 1978) and show textural evidence of metamorphism and metasomatism overprinting original magmatic cumulate textures. A small group within the suite is Type I and almost exclusively found in volcanics of basaltic andesite composition. The xenolith suite can be grouped into four components that represent different crustal levels beneath Egmont volcano (Fig. 5). 5.1. Components of the crust beneath Egmont volcano 5.1.1. Supracrustal rocks Seismic studies and drilling of oil exploration wells have showed that the uppermost 6 km of the Taranaki Basin is filled with a sequence of Tertiary sedimentary rocks. Xenoliths from these units are largely unmodified sedimentary lithologies and sub-volcanics that can be identified in either regional surface outcrops or oil well drill cores. They include Kapuni Formation quartz sandstones, Tertiary quartzofeldspathic sandstones and siltstones together with andesite volcanics from the upper part of the volcano. Fine grained sediment xenoliths show evidence of thermal alteration in both inorganic and organic constituents (Collen et al., 1985) and rarely partial melting to produce melts with up to 80% silica (Table 4, Fig. 3). 5.1.2. Upper to mid-crustal basement These comprise regionally metamorphosed and non-metamorphic basement to the Cretaceous and Tertiary sediment cover. The basement below 6 km depth is considered to comprise plutonic and

Table 5 Estimated crystallisation pressures of amphiboles in Egmont xenoliths from total number of Al cations per formula unit (based on 23 oxygen), using the geobarometer of Hollister et al. (1987). Xenolith type

Hollister et al. P (± 1 kbar)

Granodiorite Meta-hbl–pyroxenite Hbl–gabbro (green amph) Phenocrysts in andesite Hbl–px–gabbro Phenocrysts in basaltic andesite Ol–hbl–pyroxenite Hornblendite and px–hornblendite Hbl–pyroxenite (fine grained) Hbl–pyroxenite (coarse grained)

0–3.0 0–3.0 1.3–3.5 3.8–6.8 4.6–7.2 5.6–7.9 5.3–8.7 7.2–8.6 7.5–8.3 8.4–9.2

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Fig. 5. Al(IV) in hornblende from xenoliths, andesite and basaltic andesite. See text for descriptions of components 1–4. Hornblende in ultramafic xenoliths (4) has the highest Al(IV) and therefore inferred pressure while gabbroic lithologies (3) have intermediate Al(IV) or pressure and low pressures are indicated for plutonic and metamorphic lithologies (1 and 2).

metamorphic rocks of the Median Batholith (e.g. King and Thrasher, 1996; Mortimer et al., 1997). Granodiorite, granite and gneiss xenoliths are the only studied rocks that contain quartz and significant amounts of biotite and lithologies are similar to those described from the Median Batholith (Wodzicki, 1974; Challis et al., 1994; Rattenbury et al., 1998). Gneiss from the Median Batholith also contains almandine garnet, which is commonly found in regional metamorphosed sedimentary rocks, and only occurs in the gneiss rock type from the Egmont xenolith suite. Further, low Al- and Ti-contents in amphiboles from granodiorite xenoliths indicate crystallisation or recrystallisation at lower crustal pressures than most other amphibole-bearing xenoliths. This is consistent with pressure estimates from the Al in amphibole geobarometer that yields values up to 3 kbar or 10–12 km depth. The granodiorite, granite and gneiss xenoliths are therefore interpreted as fragments of basement Median Batholith rocks. 5.1.3. Mid to lower crustal cumulates and granulites This group comprises mafic hornfels, fine grained gabbro and finely banded amphibolitic gneiss xenoliths. The mineral assemblage in mafic hornfels includes wollastonite, which can form during thermal or contact metamorphism as well as in amphibolite to granulite regional metamorphism (Deer et al., 1992). Its occurrence indicates the presence of carbonate in the protolith. The granoblastic texture with randomly oriented anhedral to subhedral grains and relict minerals in the mafic hornfels is consistent with thermal rather than regional metamorphism. Clinopyroxenes are intermediate between diopside and hedenbergite, which is common for high grade metamorphosed mafic igneous rocks (Deer et al., 1992). The absence of amphibole in the mineral assemblage and the presence of relict plagioclase (An80–90) and clinopyroxene suggest an anhydrous gabbroic protolith, possibly a medium grained to fine grained cumulate gabbro that had been metasomatised by CO2-bearing fluids. It could be associated with the Median Batholith, but could also be from the Egmont volcanic–plutonic system. Fine grained gabbro found as xenoliths in basaltic andesite has essentially the same mineralogy as this mafic hornfels, apart from a slightly coarser texture and lacking relict grains, and has a similar mineral chemistry. Mafic hornfels may therefore be the more intensively metamorphosed equivalent to the fine grained gabbros. Finely banded amphibolitic gneiss xenoliths comprise mafic (clinopyroxene+ amphibole) and plagioclase (An58–75) layers. Clinopyroxenes are augites and diopsides and the latter are likely reaction products

of amphibole breakdown. Amphibole compositions encompass a similar range to those in andesites and gabbros and show evidence of decomposition to a plagioclase + clinopyroxene+ titanomagnetite assemblage. Biotite, orthopyroxene and amphibole mantling clinopyroxene indicate probable metasomatism by hydrous fluids and reaction with the high silica melts that are present (Kovács et al., 2004). One xenolith is cut by fractures that offset the mafic–felsic layering by a few millimetres. In many places, these fractures contain clear glass and euhedral pyroxene and plagioclase crystals. The presence of fractures suggests high differential stress causing brittle deformation in the lower crust that has facilitated entrainment of wall rock in magmas (O'Reilly et al., 1989). Estimated pressures from total Al in amphibole range from approximately 2.5–7.5 kbar and are comparable to results from phenocrysts in andesites, hornblende gabbros and hornblende pyroxene gabbros (Fig. 5). We interpret this petrographic type as originally cumulate gabbro in the mid to lower crust that has been subject to at least amphibolite facies metamorphism and associated deformation. Later metasomatism and reaction with siliceous partial melts formed orthopyroxene and biotite and late metasomatism also involved sulphides. 5.1.4. Gabbros and ultramafic xenoliths Hornblendite and hornblende pyroxenite xenoliths contain amphiboles and clinopyroxenes that have high Al2O3 contents. Pargasitic amphibole occurs both as a cumulus and postcumulus phase and amphibole-alteration of clinopyroxenes is also observed. Residual clinopyroxene surrounded by fine grained hornblende indicates that previously the rock was more clinopyroxene-rich and the clinopyroxene has reacted to form amphibole. Modal abundances and mineral chemistry of plagioclase and clinopyroxene in gabbroic xenoliths are comparable to those of phenocryst/xenocryst assemblages in Egmont andesites and these gabbros can be regarded as cumulates from andesite magmas that have stalled at various levels in the crust. Pressure estimates from Al in hornblende (Table 5) point to crystallisation at mid-crustal pressures. Most of these cumulate xenoliths contain evidence of hydrous metasomatism in widespread reaction of clinopyroxene to amphibole. Most amphiboles in the rock also show signs of decompression reaction to clinopyroxene and titanomagnetite plus melt which, forms glass containing fine, elongated clinopyroxene and plagioclase crystals. The meta-hornblende–pyroxenite appears to have originally been a coarse to medium grained pyroxenite or hornblende pyroxenite. Metasomatism has replaced the original clinopyroxene with fine

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grained orthopyroxene and clinopyroxene. Earlier formed amphiboles also started to break down to form clinopyroxene and titanomagnetite along felsic veins similar to those observed in finely banded amphibolitic gneiss xenoliths. For the most part the gabbros and ultramafic xenolith types are cognate cumulate xenoliths, which have also been described from other arcs, for example the Aleutians (Conrad and Kay, 1984) and the Lesser Antilles (Arculus and Wills, 1980). However, the foliated gabbroic rocks in this category may be an exception and an origin in the Median Batholith cannot be entirely excluded, as it also comprises arc-related calc–alkaline plutonic rocks. An example from a drill hole in the Tasman Sea about 70 km southwest of Mt Taranaki has been described as hornblende diorite and correlated with igneous rocks from the Rotoroa Complex (within the Median Batholith) in the South Island of New Zealand (Wodzicki, 1974). The petrographic description of this rock resembles those referred to as hornblende gabbro in this study. Hornblende gabbros found as xenoliths could therefore include mafic granulites from both earlier magmatism and earlier underplating in the Egmont magmatic system and differentiating the cumulates from granulites is difficult (Chen and Arculus, 1995; Wysoczanski et al., 1995). 5.2. Comparison with fossil arcs There are few exposures of fossil arc systems in which the complete architecture of the structures and petrological relationships of the mantle to supracrustal sequence are exposed. The most well known are the Kohistan section in northern Pakistan (Bignold et al., 2006; Garrido et al., 2006, 2007; Dhuime et al., 2007, 2009) the Talkeetna arc in south-central Alaska (Pflaker et al., 1989; DeBari and Sleep, 1991; Greene et al., 2006) Darb Zubayda in Saudi Arabia (Quick, 1990) and the Hokkaido section in Japan (Takashima et al., 2002). The two best described sections are the Kohistan Arc and the Talkeetna Arc complexes. The Cretaceous Kohistan section is interpreted as an oceanic arc that was sutured to Asia at about 100 Ma (Bignold et al., 2006; Dhuime et al., 2009). This Arc complex can be subdivided into six units (Dhuime et al., 2009). The Jijal Complex is the stratigraphically lowest unit with a lower ultramafic and upper mafic part which spans the mantle/crust boundary. It comprises dunites, wehrlites and pyroxenites grading up to garnet-bearing gabbros or “garnet granulites” with some younger granitic intrusives (Dhuime et al., 2007; Garrido et al., 2007). The Patan-Dasu metaplutonic complex comprises a thick sequence of metabasic rocks (gabbros and gabbronorites) metamorphosed and deformed under amphibolite facies conditions (Bard, 1983) and represents the main constructional phase of the oceanic arc. Lithologies represent original laccoliths interlayered with volcanic/volcaniclastic units or remnants of oceanic crust (Bard, 1983; Bignold et al., 2006). The Jaglot and Utror–Chalt meta-sediments and metavolcanics, and Yasin Group volcaniclastics, were emplaced in arc-related basins, also during the main phase of arc construction. The Chilas complex, comprising dominantly mafic (gabbronorite) intrusives represents post-suture intra-arc rifting (Garrido et al., 2006) while the Kohistan Batholith formed post-suture when the arc was continental and activity ceased when the arc was obducted during the onset of the collision between India and Asia (Garrido et al., 2006). The Early to Mid-Jurassic Talkeetna Arc accreted in the Late Jurassic to Middle Cretaceous and is thought to have developed in oceanic crust (Pflaker et al., 1989). It also comprises 6 units; a basal residual mantle harzburgite interfingering with an overlying pyroxenite unit that in turn interfingers with the succeeding basal gabbronorite. The gabbronorite is overlain by a lower crustal gabbronorite that exhibits modal layering. The upper part of the sequence comprises mid-crustal plutonics (gabbros, diorites and tonalities) succeeded by basaltic to rhyolitic supracrustal volcanics (Greene et al., 2006). In both the Talkeetna and Kohistan Arcs,

obduction has exposed relatively complete arc sections from mantle to crust of mature arc sequences. Both fossil arcs therefore exhibit a similar overall architecture of an uppermost mantle sequence of ultramafic rocks, succeeded by lower crustal gabbroic rocks, lower to mid-crustal layered gabbroic units and mid-crustal mafic to silicic intrusives. The upper part of each Arc complex comprises upper crustal volcanics. The xenolith lithologies identified from Egmont are similar to the lithologies identified from the fossil Kohistan and Talkeetna arcs. Ultramafics are represented by peridotites, pyroxenites and dunite and the Al-in-hornblende data indicate the highest pressures of formation for these rock types (Fig. 5). Lower to mid-crustal rocks are dominated by gabbroic compositions (hornblende gabbros) with hornblendites and hornblende pyroxenites. Thus the Egmont ultramafic and gabbroic xenoliths can be interpreted as representing the lithologies of the crust/mantle boundary to mid crust under the volcano. There are also superficial similarities with the fossil arcs in the presence of meta-igneous lithologies (amphibolitic gneisses, some with garnet, and siliceous plutonic rocks) but these are inherited from the older Median Batholith rocks that form the basement. Because Egmont represents the crustal section of a single volcano, isolated from the main arc in a rear arc setting of a relatively young arc, and the volcanic history spans only 0.13 Ma, the crustal section at Egmont does not exhibit the degree of arc maturity shown by the Kohistan and Talkeetna Arcs and thick volcanics and volcaniclastics are also lacking at Egmont. 5.3. Magmatic processes and crustal structure The Egmont xenolith suite represents a cross section of the subvolcanic crustal structure beneath Taranaki. Basaltic andesites have uniquely sampled depleted upper mantle (Type I xenoliths) which suggests they are sourced from deep within the magmatic plumbing system. Andesites, however, carry a xenolith cargo dominated by type II cumulates and granulites from the lower crust. Pressure estimates from Al-in-hornblende geobarometry are broadly consistent with lower pressures recorded from the Type II rocks (Fig. 5). This clearly indicates different source areas for the two magma compositions. The model of a lower crustal “hot zone “ for arc magma genesis proposed by Annen et al. (2006) is a useful concept with which to examine the processes involved. In this model underplating by basalt of the base of the crust raises the geotherm sufficiently over time to cause melting of lower crustal and previously underplated material. The basaltic andesites in the Egmont system appear to originate from the lower part of the “hot zone”, within the lithospheric mantle. In contrast the more silicic andesites were sourced from the lower crust, as reflected in the composition of their xenolith cargo. A contributing factor to the more silicic composition of the andesites is generation of silicic partial melts from the underplated lower crust. Price et al. (1999, 2005) have argued that andesites and rhyolites are formed by this mechanism and differ only in that andesites carry a large “crystal cargo” compared with the low crystal content of rhyolites. The groundmass glass compositions of Egmont andesites (Platz et al., 2007) are of similar composition to the glasses measured in the xenoliths, consistent with the Price et al. (2005) model. One outcome of these processes is progressive “andesitisation” of the middle crust as arc systems mature. 6. Conclusions Xenoliths in the Egmont suite show widespread and pervasive reaction of anhydrous minerals with fluids to form hydrous minerals such as amphibole and apatite and reaction with siliceous partial melts to form orthopyroxene. The mineralogy was further overprinted by decompression reaction of the hydrous minerals, particularly amphibole, as magmas rise to the surface.

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The Egmont xenolith lithologies tell a broadly similar story to that from mantle and lower crustal xenoliths in alkali basalts in intraplate settings of entrainment of upper mantle peridotites and lower crustal cumulates and granulites. The most significant difference is in the degree of hydrous metasomatism in the Egmont xenoliths where pervasive formation of amphibole occurs, reflecting the more fluidrich environment of arc magma systems. The overall crustal section defined by the Egmont comprises, from top to bottom, supracrustal sediments and high level cumulates, midcrustal metamorphosed basement, lower crustal cumulates and granulites and depleted upper mantle. Supracrustal xenolith lithologies are consistent with the known Tertiary sedimentary sequences in the Taranaki Basin. Type I xenoliths are only present in basaltic andesite host rocks and are sourced from depleted upper mantle, suggesting that this is also where the basaltic andesites accumulate. Type 2 xenoliths predominate in the more siliceous andesites and reflect accumulation in the lower crust. The separate source depths for the two rock types can be explained by the Annen et al. (2006) “hot zone” model, where the andesites have much greater interaction with the lower crust than the basaltic andesites. Xenolith textures, compositions and partial melt relationships are also consistent with the Price et al. (2005) model which proposes that the andesites are derived from mixing of siliceous partial melts in the lower crust with crystals residual in the source rocks. Such processes progressively “andesitise” the middle crust. Acknowledgements This paper is based on the Diplomarbeit of K. Gruender and was supported by a Johannes Gutenberg University Masters Scholarship and the DAAD (Deutscher Akademischer Austauschdienst). The authors acknowledge the long standing interest John Gamble has had in Taranaki xenoliths and express their appreciation of helpful discussions in the early stages of this project. Insightful discussions with Richard Price also helped shape this paper. We would like to acknowledge the thoughtful and constructive reviews of the manuscript by Ian Graham and an unknown referee.

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