Intraplate volcanism on the Zealandia Eocene-Early Oligocene continental shelf: the Waiareka-Deborah Volcanic Field, North Otago

ABSTRACT

Volcaniclastic deposits, pillow lavas, dikes and sills of the intraplate Waiareka-Deborah Volcanic Field in North Otago were emplaced into and onto the Zealandia Eocene-Early Oligocene continental shelf. The on-land extent is ∼890 km² but offshore volcanic rocks occurring over an additional ∼3500 km² may be related. Examination of the on-land volcaniclastic deposits indicates volcanism was dominated by short-lived clustered surtseyan-style eruptions. Pyroclast glass and sill and lava bulk rock chemistries show that the magmas were mainly sub-alkaline basalt to basaltic andesite. Minor alkaline centres are best represented by basanitic-melanephelinitic volcaniclastic deposits at Kakanui, which also contain an array of megacrysts plus mantle and crustal xenoliths. The sub-alkaline and alkaline source reservoirs had similar Sr and Nd isotope ratios (⁸⁷Sr/⁸⁶Sr_{34 Ma} = 0.70346 ± 40, ¹⁴³Nd/¹⁴⁴Nd_{34 Ma} = 0.51282 ± 4) but the sub-alkaline rocks tend to be slightly less radiogenic in ²⁰⁶Pb/²⁰⁴Pb or ²⁰⁸Pb/²⁰⁴Pb versus ²⁰⁷Pb/²⁰⁴Pb space. The dominant sub-alkaline nature and the isotopic compositions distinguish this volcanic field from the nearby Dunedin Volcanic Group and Alpine Dike Swarm. As the Waiareka-Deborah isotopic compositions are poorly represented in the Otago mantle lithosphere, the magmas may have been derived from the asthenosphere. There were multiple modes of Cenozoic intraplate volcanism in Otago.

KEYWORDS:

• Waiareka-Deborah
• submarine volcanism
• intraplate
• Zealandia

Introduction

The Waiareka-Deborah Volcanic Field near Oamaru in the South Island (Figure 1) displays spectacular evidence for submarine volcanic processes on the mid-Cenozoic Zealandia continental shelf. This intraplate volcanic field has long been known to contain locally thick piles of volcaniclastic sediments, pillow lavas, diatremes, sills and dikes (e.g. Park 1918; Uttley 1918; Benson 1943, 1944; Gage 1957; Coombs et al. 1986), including the renowned xenocryst- and xenolith-rich Kakanui Mineral Breccia (e.g. Thomson 1907; Mason 1968; Dickey 1968a; Reay and Sipiera 1987; Fulmer et al. 2010). The geochemistry of the province, however, has had little attention beyond the few analyses of Benson (1943), Coombs et al. (1986) and Hoernle et al. (2006), and regional syntheses compiled by Benson (1943, 1944) and Coombs et al. (1986) are out of date; for example, some of the alkaline intrusions formerly grouped in this province are now known to be part of the younger Dunedin Volcanic Group (Coombs et al. 2008).

Figure 1. Geological map showing the distribution of the Waiareka-Deborah Volcanic Field. The map is compiled from data presented in Benson (1943), Coombs et al. (1986), Forsyth (2001) and personal observations. Ages shown are Ar-Ar dates from Hoernle et al. (2006). Inset shows location of geological map (bold box), as well as the extent of the Alpine Dike Swarm and Dunedin Volcanic Group, which are other intraplate volcanic provinces in Otago.

We provide a synthesis of the physical volcanology and the geochemistry of the Waiareka-Deborah Volcanic Field. Phreatomagmatic deposits are spectacularly exposed in cliffs bordering the Pacific Ocean and have been well-studied in the past two decades, primarily by University of Otago students. The sparse existing crystalline rock geochemical data are supplemented with new major and trace element and Sr-Nd-Pb isotopic analyses to give a regional perspective of the magmatism and its mantle sources. These data enable this intraplate province to be compared with the spatially overlapping but younger (25-9 Ma) Dunedin Volcanic Group as well as the younger and spatially distinct (∼25-20 Ma) Alpine Dike Swarm, both of which are also intraplate volcanic fields in the Otago region.

Geological setting

After breakaway from Gondwana at 84 Ma, New Zealand’s sedimentary record shows this continental mass became largely submerged. The peak of marine transgression occurred in the Early Oligocene (e.g. Forsyth 2001), with later marine regression and uplift of the North and South Islands initiated by mid-Oligocene propagation of the Australia-Pacific plate boundary through the centre of the continent. Zealandia has throughout its history been punctuated by volcanism. While the most prominent volcanism, forming the volcanoes of the Taupo Volcanic Zone, is arc-related, Zealandia has a long history of intraplate volcanism (e.g. Mortimer and Scott 2020) and one product of this history is the topic of this paper: mid-Cenozoic volcanic rocks near Oamaru in North Otago in the South Island. The volcanic rocks around Oamaru were historically considered to comprise two separate formations with the older Waiareka Volcanic Formation separated from the younger Deborah Volcanic Formation by the Totara Limestone (e.g. Park 1918; Uttley 1918, 1920; Benson 1943; Coombs and Dickey 1965). This limestone, subsequently renamed as Ototara, was deposited during and between times of volcanic activity but the inference that there is an older and a younger member separated by a hiatus involving regional carbonate formation is a major oversimplification. Instead, there were many local periods of non-volcanic deposition that produced limestone and marine mud- and silt-stone units (Coombs et al. 1986; Edwards 1991; Andrews 2003; Maicher 2003). The two volcanic formations are now grouped as part of the Alma Group (Forsyth 2001) and following Coombs et al. (1986) we refer to the volcanic rocks collectively as the Waiareka-Deborah Volcanic Field.

The minimum extent of the Waiareka-Deborah Volcanic Field is ∼890 km² (Figure 1). However, erosion has likely removed some inland occurrences and there is evidence in seismic lines and specimens from test wells that indicates intrusions and tephra offshore (Wilding and Sweetman 1971; Field et al. 1989; Bischoff et al. 2020). If the offshore volcanism is included within the Waiareka-Deborah Volcanic Field (∼3,500 km; Bischoff et al. 2020), then the entire volcanic field occurs over an area probably closer to ∼4400 km². Biostratigraphic ages from units bounding tuffs from Endeavour-1 Well indicate, however, that this areal extent also includes Paleocene volcanic rocks (Wilding and Sweetman 1971), and so the extent of offshore volcanism correlating with the on-land Waiareka-Deborah Volcanic Field remains unknown.

The crustal ‘basement’ underlying the Waiareka-Deborah Volcanic Field is the dominantly quartzofeldspathic Otago Schist, which was metamorphosed to greenschist facies conditions in the Late Jurassic – Early Cretaceous and exhumed by ∼ 100 Ma (Mortimer 2000; Forsyth 2001). However, felsic granulite facies xenoliths in the Kakanui Mineral Breccia indicate that the middle and lower crust, at least in part, comprises rocks metamorphosed at ∼92 Ma under extreme high to ultra-high temperature conditions (∼900°C; Jacob et al. 2017). The Otago Schist is unconformably overlain in the study area by terrestrial Cretaceous quartzose conglomerate, sandstone and mudstone belonging to the Matakea Group, and/or Late Cretaceous to Early Oligocene Onekakara Group sediments that evolve from non-marine to marine (Forsyth 2001). The volcanic rocks of the Alma Group – the topic of our paper – were emplaced during the Eocene-Late Oligocene, very late in the accumulation of region-wide Onekakara Group sandstone and mudstone (e.g. Thompson et al. 2014). A regional unconformity, the so-called ‘Marshall Paraconformity’, separates the Alma Group and Onekakara Group from Late Oligocene to Miocene Kekendon Group sediments. It was during Kekendon Group sedimentation that a significant portion of Zealandia emerged above sea-level and this was accompanied by Dunedin Volcanic Group magmatism, which is an extensive alkaline volcanic province that in part overlaps spatially with the Waiareka-Deborah volcanism (Figure 1) (Coombs et al. 1986, 2008; Scott et al. 2020). Emplacement of the intraplate alkaline-carbonatitic Alpine Dike Swarm in West Otago also occurred in the Oligocene to Early Miocene (e.g. Cooper 2020).

Volcaniclastic rocks

Volcaniclastic rocks are the most areally extensive component of the on-land Waiareka-Deborah Volcanic Field (Figure 1), although they are very poorly exposed except in coastal sections. Detailed sedimentological descriptions of these deposits have been made, from north to south, from Cape Wanbrow (Uttley 1918; Coombs et al. 1986; Moorhouse et al. 2015; Moorhouse and White 2016), the areas around Maheno and Clarks Mill (Hicks 2014), Kakanui (Dickey 1968a; Corcoran and Moore 2008), Aorere Point-Bridge Point (Cas et al. 1989; Hicks 2014), Lookout Bluff (Maicher 2003) and Moeraki (Andrews 2003) (Figure 1). These sequences are more than 150 m thick in some places (Coombs et al. 1986; Andrews 2003; Corcoran and Moore 2008; Moorhouse et al. 2015). A road cutting now largely concreted over once exposed a few-metre thick section of mantle xenolith-bearing tuff breccia and lapilli tuff overlying pillow basalt at Alma (Reay et al. 2002).

All of the studied sections comprise sequences of tuff breccia, lapilli tuff and tuff (Figure 2A) that are in places interbedded with pillow lava and limestone, mudstone or marl (Figure 2B). The contact relationships among different volcaniclastic deposits, often featuring sharp but local erosion surfaces (Figure 2C) and/or cross-bedding, indicate that they formed primarily as deposits of submarine density currents from proximal volcanic edifices, emplaced both as primary deposits during eruptions, and through subsequent post-eruptive redistribution (Cas et al. 1989; Andrews 2003; Maicher 2003; Corcoran and Moore 2008; Moorhouse et al. 2015). Volcaniclasts are mostly now-altered vesicular basaltic glass, or blocks fragmented from dikes associated with the eruptions (Figure 2D). Anomalously, Kakanui deposits also contain a variety of crust and mantle fragments (described below). Palaeontology indicates that water depths were generally shallow (∼ <80 m) during the volcanic field's duration (Lee et al. 1997; Hicks 2014) and it is probable that volcanic edifices breached the sea surface, with those of entirely pyroclastic construction being rapidly planed off to normal wave base (e.g. Thorarinsson 1967). The volcanoes eroded rapidly, with evidence for calcareous algae, bryozoans and serpulids encrusting basalt boulders and cobbles (Lee et al. 1997) indicating the volcanoes formed sites favourable for rhodolith, bryozoan, brachiopod, echinoderm, bivalve and gastropod colonies (Utlley 1918; Cas et al. 1989; Lee et al. 1997; Corcoran and Moore 2008; Hicks 2014).

Figure 2. Examples of Waiareka-Deborah Volcanic Field phreatomagmatic rocks. A, View of fine-grained tephra beds that are extensive around Cape Wanbrow. The cliffs are ∼30 m high. B, A spectacular sequence of tephras and thin calcareous tuffs, tuffaceous limestone and limestone horizons is exposed at Boatmans Harbour on the northern side of Cape Wanbrow. The occurrence of limestone beds, and elsewhere limestone or mudstone, indicates that the volcanic field had numerous hiatuses between intermittent eruptions. Photo courtesy of Marco Brenna. C, Truncated bedforms are very common at Kakanui and the erosion surfaces are of only local significance. D, Breccia components typical comprise basalt, in places appearing to be made from fragmented dikes. Outcrop is at Lookout Bluff. Pocket knife is ∼8 cm long.

It is widely accepted that the Waiareka-Deborah volcaniclastic deposits are products of phreatomagmatic interaction of magma with seawater, which drove surtseyan-style eruptions on the continental shelf (Figure 3) (Coombs et al. 1986; Cas et al. 1989; Andrews 2003; Maicher 2003; Corcoran and Moore 2008; Moorhouse et al. 2015). The volcaniclastic sedimentary record shows that in multiple sites deposits from multiple volcanoes, formed at different times, overlap in the stratigraphy (Maicher 2003; Moorhouse et al. 2015). The presence of limestone, mudstone and glauconitic siltstone between volcaniclastic deposits shows, however, that there were significant hiatuses between the eruptions that formed different small volcanoes in the field, even where they are spatially clustered (Andrews 2003; Maicher 2003; Corcoran and Moore 2008). Formation of tephra deposits in the basin away from immediate volcano flanks was probably by some combination of (1) distal eruption-fed density currents (e.g. Verolino et al. 2018); (2) vertical density currents driven by accumulation of ash near the water surface (e.g. Bradley 1965; Jacobs et al. 2015), and; (3) settling of ash through the water column (e.g. Kutterolf et al. 2018).

Figure 3. Cartoon showing the inferred shallow-level development of the Waiareka-Deborah Volcanic Field.

Original pyroclasts preserved in the volcaniclastic rocks are commonly extensively palagonitised and only plagioclase crystals survive with any regularity (e.g. Moorhouse et al. 2015). Nonetheless, sideromelane with small crystals of olivine, plagioclase and clinopyroxene has been found dispersed through the rocks (Coombs et al. 1986; Andrews 2003), with the glass chemistry indicating mostly transitional alkaline to sub-alkaline compositions (Coombs et al. 1986). Evidence for alkaline eruptions is recorded from a basanitic ash horizon at Cape Wanbrow and sideromelane pyroclasts at Kakanui (Dickey 1968a; Coombs et al. 1986). The volcaniclastic record is therefore interpreted to show that were there numerous small surtseyan-style eruptions in the field (Figure 3) with different magma batches for each volcano.

Lavas, sills and dikes

Pillow lavas, sills and dikes make up a smaller part of the outcrop area than do volcaniclastic rocks in the on-land portion of the volcanic field (Figure 1), and as non-fragmental deposits they provide information about non-explosive eruptions, or parts of eruptions. The most comprehensive descriptions of these coherent rocks, although reconnaissance in nature, remain those made by Uttley (1918), Benson (1943) and Nakamura and Coombs (1973). Sills and intrusions have also been mapped in offshore seismic sections (Bischoff et al. 2020), but none have been sampled or dated.

The best exposures of pillow lava occur at Boatmans Harbour, where a 30 m thick horizon of basaltic pillows set within Ototara Limestone matrix is spectacularly exposed (Park 1918; Uttley 1918; Kawachi and Pringle 1988) (Figure 4A, B). The pillow rinds are unaltered sideromelane enclosing fresh magmatic labradorite and olivine, with clinopyroxene present in the pillow interiors (Coombs and Dickey 1965; Coombs et al. 1986; Kawachi and Pringle 1988). Single pillows commonly have more than one rind. This feature may indicate fragmentation of initial rinds followed by water invasion and formation of new ‘interior’ rinds, either as pillows partly imploded during emplacement (Kawachi and Pringle 1988) or as rinds were broken and partly detached during pillow expansion (Walker 1992). Slightly higher in the sequence, and just above an intervening thick limestone horizon, a thick tuff breccia represents a lava-fed delta, containing broken pillows and other glassy fragments (Figure 4B). Lava in Awamoa Creek reportedly comprises pillows up to ‘18 feet in diameter’ (Benson 1943, p. 120). Very altered pillow lava occurs at the Alma road cut (Reay et al. 2002) and Oamaru Creek (Uttley 1918) but these have not been examined in much detail.

Figure 4. Flows, sills and dikes of the volcanic field. A, Pillow lava overlain by tuffaceous material at Boatmans Harbour (Cape Wanbrow; sequence also illustrated in Figure 2B). The beach to the top of the cliffs on the right-hand side of 4A are about 15 m high. B, The pillows have glassy margins and have interstitial bryozoan limestone. C, The Tokorahi Sill, here exposed on Dip Hill Road, has sub-vertical columnar jointing. The basal contact with Eocene sandstone is just to the left of this photo. Although not obviously differentiated at outcrop scale, the whole rock geochemistry shows come variation in composition across the body. D, View of the Mt Charles Sill, which intrudes concretion-bearing mudstone. The dolerite is extremely friable on this cliff, but fresher (although less well exposed) material occurs at the southern end of the sill. This sill is chemically differentiated vertically, from olivine dolerite at the base upwards to quartz dolerite (Benson 1943). View is to the northeast from State Highway 1. The exposed sill is ∼20 m thick. E, The Tawhiroko Sill, which is also geochemically differentiated (Benson 1943), on the Moeraki Peninsula has horizons rich in schist clasts that are derived from the underlying Otago Schist. Photos courtesy of M Brenna. F, Features of mafic dikes on Moeraki Peninsula. The dikes appear to be made of multiple injections and are packed, in places, with calcite-filled vesicles. Inset shows the ∼6 cm wide white porcellanite margin to one dike. Photo courtesy of AF Cooper.

Doleritic sills, emplaced at very shallow levels, are a widespread feature of the volcanic field (Figures 1 and 3). The sills intrude Eocene sandstone at Tokarahi near Maerewhenua, limestone at Clarks Mill near Maheno, mudstone at Mt Charles, and siltstone and mudstone on the Moeraki Peninsula (Benson 1943, 1944) (Figure 1). Dolerite at Round Hill may also be a sill (Benson 1943), although this has not been confirmed. The most regionally extensive of these sills are the ∼20 m thick columnar-jointed intrusion at Tokarahi (Figure 4C), which may have been emplaced at the same time as Basalt Hill to the north (Figure 1), and the ∼ 50 m thick Mt Charles Sill (Figure 4D). Each of these sills is estimated to have an extent of over ∼25 km² and volumes on the order of 1–2 km³ (Coombs et al. 1986). The Mt Charles Sill is underlain by flaggy concretion-bearing mudstone with the sill itself grading upwards from (weathered) olivine dolerite into quartz dolerite (Benson 1943; Coombs et al. 1986). The best studied sills, however, are those on the Moeraki Peninsula. Here, the Tawhiroko Sill is at least 50 m thick (the top having been lost to erosion) and has olivine dolerite at the base but a pegmatitic quartz dolerite core (Benson 1943, 1944). It also contains horizons remarkably rich in schist xenoliths, locally size-graded in bands within the sill (Figure 4E) despite Otago Schist being 10–100 m below the sill. Nakamura and Coombs (1973) interpreted chemical zoning in clinopyroxene grains in this sill to be due to progressive crystallisation of the magma. Benson (1943) attributed the chemical differentiation in the Waiareka-Deborah sills to gravitational settling, but the differentiation could also be due to emplacement of variably evolved magmas injected as different batches. This would better explain the occurrence of schist-rich horizons at Moeraki. Coombs and Roedder (1994) reported the occurrence of CO₂ inclusions in plagioclase microphenocrysts from some Moeraki Peninsula sill rocks and inferred these to represent immiscible CO₂ droplets in the melt at ∼ 10 km depth.

Doleritic dikes are widespread but exposures are concentrated on Moeraki Peninsula and around Enfield (Benson 1943; Coombs et al. 1986; Andrews 2003). Dikes on Moeraki Peninsula are mostly < 1 m wide, but range up to several metres width for those extending in outcrop for many 10s of metres. Dikes characteristically show evidence for multiple phases of injection (Figure 4F). Marginal porcellanite from one Moeraki Peninsula dike was apparently quarried for tools by Māori prior to the arrival of Europeans (Benson 1943). A complication is that some dikes cropping out in the area mostly occupied by the Waiareka-Deborah Volcanic Field may belong to the younger alkaline Dunedin Volcanic Group (Coombs et al. 2008; Scott et al. 2020); chemical and isotopic analysis is required to test dike heritage.

The mantle under the Waiareka-Deborah Volcanic Field

Mantle xenolithic material occurs in five locations (Scott 2020). Very small (∼1 cm) spinel peridotite fragments occur in tuff beds above the intact pillow lava at Boatmans Harbour (Coombs et al. 1986; Moorhouse 2015) (Figure 4A). The second occurrence is in lapilli tuff ∼ 2.5 km east of Round Hill, which is packed with moderately altered spinel peridotite xenoliths up to ∼ 8 cm in diameter (Scott et al. 2014b). A third occurrence is in thin lapilli tuff and tuff breccia beds directly beneath Ototara Limestone at Alma. These beds contain spinel peridotite and garnet pyroxenite xenoliths, as well as garnet, kaersutite, anorthoclase and augite megacrysts (Reay et al. 2002; Scott et al. 2014b). The peridotite clasts from Alma are extremely altered, with the mantle orthopyroxene and olivine having been completely converted to clay (Scott et al. 2014b). It is not known whether spinel lherzolite-bearing basaltic boulders in the river at Five Forks originate from the Dunedin Volcanic Group or Waiareka-Deborah-Volcanic Field.

In contrast to the aforementioned xenolith occurrences, the intra-vent lapilli tuff breccia enclosed within bedded surtseyan lapilli tuff on a marine platform at Kakanui (Figure 5A) contains a spectacular array of xenocrystic and xenolithic material set in a calcite and zeolite-cemented primary volcaniclastic matrix (Figure 5B) (Mantell 1850; Dickey 1968a, Dickey 1968b; Reay and Sipiera 1987; White and Houghton, 2006). The most xenolith-rich location, South Head, is the centre of a single surtseyan volcano (Corcoran and Moore 2008). Here, the Kakanui Mineral Breccia contains spinel lherzolite and harzburgite, cut in places by amphibole-phlogopite-bearing veins, as xenoliths up to about 20 cm in diameter (but commonly much smaller and usually heavily replaced by carbonate) within thin-rinded bombs of melanephelinite (Figure 5C, D) (Dickey 1968a; Reay and Sipiera 1987; Klemme 2004; McCoy-West et al. 2013; Scott et al. 2014b). These xenoliths are fragments of the mantle lithosphere from under the volcanic field. Olivine (Mg# = 100*Mg/(Mg + Fe) = 89.3–91.0) and spinel (Cr# = 100*Cr/(Cr + Al) = 8.8–44.4) compositions from lherzolites and harzburgites indicate that this mantle is fairly to moderately fertile (Reay and Sipiera 1987; McCoy-West et al. 2013; Scott et al. 2014b; Moorhouse 2015). A single whole rock Os analysis (¹⁸⁷Os/¹⁸⁸Os = 0.11953; McCoy-West et al. 2013) and three clinopyroxene Hf analyses (εHf = +100 to +14; Scott et al. 2014b) indicate depletion ages that range back to Proterozoic time; the geological significance of the old ages remains debated, with suggestions that the underlying mantle is either a vast block of ancient lithosphere (McCoy-West et al. 2013) or that it is young lithosphere with embedded ancient fragments (Liu et al. 2015; Scott et al. 2019). Enriched clinopyroxene trace element data show that the underlying mantle was metasomatized after depletion (Scott et al. 2014b; McCoy-West et al. 2015), and this is supported by the occurrence of apatite crystals and crystallised fluoride melt in one xenolith (Klemme 2004) and the amphibole-phlogopite-bearing veins (Figure 5D).

Figure 5. Kakanui Mineral Breccia outcrops and hand specimens. A, Aerial view of the Kakanui Mineral Breccia, located near the centre of a submarine diatreme. The breccia is overlain by loess and obscured by the Kakanui township. Image from Google Earth. White arcuate lines indicate bedding orientations. B, The tuff breccia facies contains xenoliths of lherzolite (commonly carbonated), garnet pyroxenite, nephelinite with amphibole megacrysts/xenocrysts, all set within a calcite-zeolite cement. Fine-grained dark clasts are basaltic/melanephelinite fragments. C, Rare peridotite xenoliths up to about 20 cm in size have a thin nephelinite annulus. Photo courtesy of DG Pearson. D, Peridotite xenoliths have cross-cutting amphibole veinlets. E, Megacrysts form an important component to the mineral breccia. The most common are amphibole, followed by anorthoclase. Garnet is rare. Many grains have a rounded nature with a polished surface, possibly due to abrasion during entrainment and emplacement. F, View of granulite xenoliths. The xenoliths provide insight into the middle-lower crust under the region. Felsic granulites are most common, and have been found to preserve evidence for Late Cretaceous ultra-high temperature metamorphism of the lower Otago crust.

Clinopyroxenite and garnet pyroxenite are common in the Kakanui Mineral Breccia (Figure 5B) (Mason 1968; White et al. 1972; Reay and Sipiera 1987; Zack et al. 1997; Sun 2018). The clinopyroxenites have attracted little attention. However, the garnet pyroxenite xenoliths, which were considered to be eclogite but lack sufficiently aluminous clinopyroxene for this classification, are mainly picro-basalt to basaltic in composition (Figure 6). Metamorphic equilibration temperatures are in excess of 950°C (Zack et al. 1997), and some record Cretaceous Lu-Hf metamorphic ages (Scott, unpublished data). This latter result means that the garnet pyroxenites are not related to Eocene-Oligocene magmatism in the field.

Figure 6. Total alkali versus SiO₂ diagram showing data from Cenozoic intraplate volcanic provinces in Otago. Waiareka-Deborah lava and sill data are presented in Table 1; the Dunedin Volcanic Group data are from the compilation presented by Scott et al. (2020); Alpine Dike Swarm data are from the compilation of Cooper (2020); and the Kakanui garnet pyroxenite xenolith data are unpublished.

The tuff breccia at Kakanui contains a remarkable array of xenocrysts. These are, in order of estimated decreasing occurrence: kaersutitic to pargasitic amphibole, anorthoclase, augite, pyrope (Figure 5E), as well as biotite, ilmenite and apatite (Dickey 1968a; Reay and Wood 1974; Merrill and Wyllie 1975; Wallace 1977; Reay and Sipiera 1987; Reay et al. 1989, 1993; Fulmer et al. 2010; Urosevic et al. 2018). Many of the xenocryst margins are curved and shiny in outcrop, and many are enclosed in melanephelinite lapilli. Several large crystals of kaersutite, pyrope, anorthoclase and augite have been distributed globally as chemical micro-analytical standards and are available from the Smithsonian Institute upon request. Whilst it is reported that some megacrysts reach 20 cm in length (Reay and Sipiera 1987), exposed grains of this size have long been collected and only the amphiboles, anorthoclase and augite of up to about 5 cm diameter can now be readily found. Lu-Hf isotope data indicate that pyrope was in equilibrium with the melanephelinite, whereas the amphibole megacrysts were not (Fulmer et al. 2010); we have no comparable information for other megacryst species.

An additional feature of the Kakanui Mineral Breccia is that it contains abundant granulite xenoliths. These are geodynamically significant for Otago because they preserve evidence for an ultra-high-temperature metamorphic event (∼900°C) that affected the base of the Otago crust at ∼92 Ma (Jacob et al. 2017), an event for which there is no known evidence at the Otago surface. Among the granulite xenoliths, weakly foliated felsic ones are common (Figure 5F) and mafic garnet-pyroxene granulites are rare. Jacob et al. (2017) interpreted the granulites to have formed by bottom-up heating of the Otago crust caused by Late Cretaceous slab rollback coupled with concomitant regional crustal extension.

Age

Age determinations for Waiareka-Deborah volcanism have historically been established by paleontological means because deposits are enclosed in strata containing molluscs, brachiopods and foraminifera. These indicate that volcanism mostly occurred in the New Zealand’s Eocene Runangan (36.4–34.6 Ma) and Late Eocene to Early Oligocene Whaingaroan (34.6–27.6 Ma) stages (Benson 1943; Gage 1957; Coombs et al. 1986; Hicks 2014). Published radiometric dates are restricted to ⁴⁰Ar/³⁹Ar analyses of 33.6 ± 1.8 Ma (Round Hill), 34.0 ± 0.6 Ma (Maheno) and 34.2 ± 0.4 Ma and 34.3 ± 0.9 Ma (Boatmans Harbour) (Hoernle et al. 2006) (Figure 1), all of which confirm Runangan to Early Whaingaroan emplacement. Two basaltic clasts in tuff breccia at Bridge Point gave ages of 39.5 ± 1.8 Ma (Eocene Bortonian or Kaiatan) and 34.3 ± 0.5 Ma (Runangan) (Hoernle et al. 2006); the older age incorporates 69% of the ³⁹Ar and is therefore probably significant despite immediately overlying fossiliferous volcaniclastic deposits being Runangan in age (Lee et al. 1997; Hicks 2014). Two fractions of the same megacrystic kaersutite from Kakanui, KK1, prepared by Reay et al. (1989), yielded Runangan-Whaingaroan ages of 33.7 ± 0.3 and 34.1 ± 0.1 Ma (Hoernle et al. 2006). Fulmer et al. (2010) report an approximate isochron age of 34 Ma for garnet, amphibole and whole-rock melanephelinite from Kakanui. In summary, palaeontological and radiometric data confirm that most Waiareka-Deborah volcanism occurred over a geologically short duration that definitely initiated in the Eocene and probably extended into the Early Oligocene.

Geochemistry

There have been few whole-rock geochemical studies of the Waiareka-Deborah Volcanic Field because most volcaniclastic deposits are altered (Coombs et al. 1986; Moorhouse et al. 2015). We supplement the existing major and trace data of Hoernle et al. (2006; n = 3), Hoke et al. (2000; n = 1), Klemme (2004; n = 1), Reay and Sipiera (1987; who reported the average of six clasts), and trace elements for one sample (Fulmer et al. 2010), with a further 9 analyses from sills and flows sampled across the volcanic field (Table 1). The whole-rock data were obtained from ALS Minerals in Brisbane on fused Li-borate disks, following the method reported by Scott et al. (2020). Although these data enable characterisation of the sampled coherent volcanic rocks, their volume is far exceeded by fragmental rocks; results may not be representative for the field as a whole. The samples are also commonly slightly altered, as indicated by LOI values reaching up to 5.5 wt%; there has probably been loss of some mobile elements.

Table 1. New and published whole rock geochemical and isotopic data. Oxides are in weight %; trace elements are in ppm; locations are latitude, longitude.

Location	Tawhiroko Sill	Moeraki Sill	Clarks Mill Sill	Clarks Mill Sill	Tokarahi Sill	Tokarahi Sill	Tokarahi Sill	Mt Charles Sill
Sample	52319	5872	55062	5721	Tok-1	Tok-2	5725	5727
Location	−45.371, 170.866	−45.356, 170.857	−45.158, 170.851	−45.158, 170.851	−45.940, 170.618	−44.925, 170.618	uncertain	−45.254, 170.799
Data Source	This study	This study	This study	This study	This study	This study	This study	This study
SiO2	51.60	52.10	51.70	51.70	49.20	52.20	54.30	50.90
TiO2	2.45	2.02	1.84	1.81	1.58	1.73	1.82	2.03
Al2O3	12.80	15.00	14.45	14.60	14.00	14.40	15.65	13.65
Fe2O3	11.30	10.25	11.40	11.05	10.45	11.05	7.40	11.30
MnO	0.14	0.15	0.13	1.14	0.10	0.12	0.08	0.14
MgO	4.03	3.71	7.13	7.15	7.32	6.73	2.91	4.25
CaO	6.44	8.47	8.25	8.23	8.57	8.68	9.46	8.40
Na2O	4.16	3.61	3.76	3.77	2.93	3.27	3.43	3.85
K2O	1.00	0.89	0.82	0.81	0.27	0.61	0.54	0.75
P2O5	0.38	0.33	0.30	0.30	0.22	0.23	0.22	0.27
LOI	5.55	3.80	0.43	0.30	1.99	1.20	3.68	4.72
Total	99.85	100.33	100.21	100.86	96.63	100.22	99.49	100.26
Mg#	41	42	55	56	58	55	44	43
Cs	0.39	0.24	0.17	0.19	0.25	0.59	0.62	0.34
Sr	416	541	444	437	406	420	465	400
Rb	25.9	10.0	16.1	17.5	2.2	11.9	6.2	17.1
Ba	258	221	180.5	173.5	130	143	167.5	170
Th	4.0	2.1	2.9	2.9	1.8	1.9	2.1	2.2
U
Nb	29.9	27.1	25.7	25.1	15.6	17	18.3	18.7
Ta	1.8	1.7	1.7	1.6	1.1	1.2	1.2	1.3
La	28.4	23.4	17.5	17	12	12.7	13	15.8
Ce	58.5	47.1	37.1	35.8	25.7	27.5	28.6	33.8
Pb
Pr	6.54	5.53	4.41	4.36	3.09	3.42	3.35	4.02
Nd	28.3	24.3	18.8	18.9	13.9	15.6	15.6	18.3
Zr	202	167	162	172	111	124	122	138
Hf	5.1	4.2	4	4.3	2.9	3.1	3.3	3.5
Sm	7.1	6.2	4.7	4.6	3.7	4.2	3.9	4.7
Eu	2.3	2.0	1.7	1.7	1.4	1.6	1.6	1.8
Gd	7.5	6.1	4.9	5.0	4.2	4.8	4.7	5.6
Tb	1.08	0.9	0.75	0.74	0.64	0.71	0.70	0.83
Ga	24.7	22.9	21.8	21.9	19.2	20.6	21.8	22.5
V	202	178	174	170	156	166	170	205
Dy	6.3	5.1	4.0	3.9	3.5	3.8	3.7	4.6
Ho	1.07	0.92	0.78	0.77	0.65	0.71	0.73	0.85
Er	3.04	2.45	2.09	2.06	1.68	1.9	1.86	2.16
Tm	0.37	0.31	0.26	0.26	0.22	0.25	0.26	0.31
Y	29.7	24.2	20	19.7	17.5	19	19	22.8
Yb	2.23	1.77	1.54	1.44	1.29	1.48	1.45	1.72
Lu	0.36	0.19	0.24	0.24	0.19	0.24	0.22	0.22
Cr	20	110	280	280	300	290	310	60
Ni
^87Sr/^86Sr	0.70362	0.70352	0.70345	0.70346	0.70368	0.70354	0.70387
2 sigma	0.00001	0.00001	0.00001	0.00001	0.00001	0.00001	0.00001
^87Rb/^86Sr	0.174	0.053	0.105	0.116	0.016	0.082	0.039
^87Sr/^86Sr(i)	0.7035	0.70349	0.70345	0.70340	0.70368	0.70350	0.70385
^143Nd/^144Nd	0.51285	0.51288	0.51285	0.51284	0.51288	0.51287	0.51286
2 sigma	0.00002	0.00001	0.00002	0.00001	0.00001	0.00001	0.00001
^147Sm/^144Nd	0.151	0.153	0.150	0.147	0.160	0.161	0.151
^143Nd/^144Nd(i)	0.51282	0.51288	0.51282	0.51281	0.51285	0.51283	0.51283
^206Pb/^204Pb	19.121	19.158	19.155	19.183	19.109	19.125	19.085
2 sigma	0.001	0.001	0.001	0.001	0.001	0.001	0.001
^207Pb/^204Pb	15.642	15.640	15.648	15.647	15.649	15.649	15.648
2 sigma	0.001	0.001	0.001	0.001	0.001	0.001	0.001
^208Pb/^204Pb	38.853	38.864	38.883	38.921	38.834	38.848	38.819
2 sigma	0.004	0.003	0.004	0.003	0.003	0.003	0.003

Location	Awamoa Pillow Lava	Boatmans Harbour Pillow Lava	Trig S, Maheno	Round Hill	Kakanui	Kakanui	Kakanui
Sample	5753	OU47055	OU54929	OU55008	OU53628	OU53628	Av. 6 clasts
	uncertain	−45.113, 170.983	−45.175, 170.842	−45.102, 170.851	−45.193, 170.902	−45.193, 170.902	−45.193, 170.902
Data Source	This study	Hoernle et al.	Hoernle et al.	Hoernle et al.	Hoernle et al./ This study	Hoke et al.	Reay & Sipiera
SiO2	48.20	51.07	49.43	48.84	38.31	39.09	44.49
TiO2	1.70	1.78	1.80	1.67	2.77	2.92	2.86
Al2O3	14.30	14.79	14.73	14.18	12.20	12.50	18.11
Fe2O3	11.25	11.46	11.64	11.83	14.59	15.31	11.53
MnO	0.12	0.14	0.15	0.15	0.20	0.22	0.02
MgO	7.81	7.78	7.49	7.93	9.81	9.93	3.84
CaO	8.15	8.23	8.17	8.01	11.14	12.65	9.93
Na2O	3.16	3.37	3.57	3.34	3.43	3.70	4.53
K2O	0.29	0.80	0.99	0.67	1.50	1.52	3.18
P2O5	0.23	0.28	0.44	0.26	1.55	1.64
LOI	2.03	0.71	1.61	3.06	5.42
Total	97.24	100.41	100.02	99.94	100.92	99.48	98.49
Mg#	58	57	56	57	57	56	40
Cs	0.34	0.46	0.41	0.13	0.63
Sr	384	438	634	493	1378	83
Rb	4.1	18	27	16	26	26
Ba	112	173	234	151	580	623
Th	2.1	2.5	3.5	2.1	11.9
U		0.74	1.00	0.57	1.75
Nb	17.8	21.3	36.1	17.9	106.6	108.0
Ta	1.1	1.3	1.7	0.8	5.4	5.7
La	12	13.9	23.8	12.7	101.8	100.0
Ce	27.3	29.4	48.1	26.9	192.2	189.7
Pb		2.2	2.9	2.3	5.3
Pr	3.2	3.81	5.68	3.35	22.56
Nd	15.0	16.4	22.8	14.3	85.1	83.2
Zr	112	129	169	115	297	300
Hf	2.9	3.2	3.1	2.4	6.3	6.3
Sm	3.7	4.2	5.0	3.6	14.8	18.97
Eu	1.5	1.5	1.7	1.3	4.5	4.64
Gd	4.3	4.5	4.5	3.5	12.3
Tb	0.62	0.69	0.68	0.57	1.55	1.31
Ga	19.9	20.2			21.1	21.00
V	167	170	224	209	194	180
Dy	3.6	3.8	3.6	3.1	7.1
Ho	0.65	0.70	0.63	0.58	1.12
Er	1.87	1.73	1.56	1.45	2.46
Tm	0.22	0.23	0.20	0.19	0.28
Y	17.9	17.5	21.3	19.6	27.3	32.0
Yb	1.28	1.44	1.19	1.20	1.52	1.32
Lu	0.25	0.20	0.16	0.17	0.19	0.15
Cr	260	260	319	384	115	111
Ni		165	188	241	145	166
^87Sr/^86Sr	0.70347	0.70327	0.70321	0.70342	0.70342
2 sigma	0.00001	0.00001	0.00001	0.00001	0.00001
^87Rb/^86Sr	0.031	0.119	0.123	0.094	0.055
^87Sr/^86Sr(i)	0.70346	0.70321	0.70310	0.70337	0.70339
^143Nd/^144Nd	0.51286	0.51284	0.51286	0.51284	0.51290
2 sigma	0.00001	0.00001	0.00001	0.00001	0.00001
^147Sm/^1^44Nd	0.149	0.156	0.133	0.153	0.105
^143Nd/^144Nd(i)	0.51283	0.51281	0.51283	0.51281	0.51288
^206Pb/^204Pb	19.042	19.118	19.247	19.099	19.313
2 sigma	0.001	0.001	0.003	0.002	0.001
^207Pb/^204Pb	15.642	15.617	15.631	15.637	15.622
2 sigma	0.001	0.001	0.003	0.002	0.001
^208Pb/^204Pb	38.798	38.763	38.871	38.792	38.948
2 sigma	0.003	0.002	0.006	0.004	0.003

Using the total alkali versus silicate classification (TAS: Le Maitre 2002), the Waiareka-Deborah whole-rock compositions of coherent volcanic rocks (lavas, dikes or sills) are sub-alkaline basalt to basaltic andesite, except for the alkaline melanephelinite and basanitic clasts in the Kakanui Mineral Breccia (Figure 6). The occurrence of both alkaline and sub-alkaline components is consistent with the chemistry of glass preserved in volcaniclastic deposits (Coombs et al. 1986). The Mg# (=100*Mg/(Mg + Fe)) of the sub-alkaline volcanic rocks ranges from 42 to 60, which is very similar to the Kakanui melanephelinite (57-56) but higher than the Kakanui basanites (40-33) (Table 1). In samples not belonging to the sills that are clearly chemically differentiated, the high MgO (>7 wt%), Cr (∼300 ppm) and Ni (>250 ppm) indicates that these rocks are fairly primitive. Trace element data from the two Kakanui analyses show the rock is slightly to significantly more enriched incompatible element abundances than are the sub-alkaline rocks (Figure 7A, B). La/Lu_(N) values for the sub-alkaline components (5.1–15.8) are less than for the alkaline Kakanui components (20.1–57.7). Some sub-alkaline samples display negative K and Rb anomalies, and all sub-alkaline rocks display positive Sr and P anomalies.

Figure 7. A, Trace element data for the Waiareka-Deborah Volcanic Field, compared to summaries of mafic magmas in the nearby Dunedin Volcanic Group (Scott et al. 2020) and Alpine Dike Swarm (Cooper 2020). B, REE data for Waiareka-Deborah sub-alkaline and alkaline magmas. Normalising values in both diagrams are from Sun and McDonough (1989). Otago Schist greyschist composition is from Scanlan et al. (2020).

Sr-Nd-Pb isotopes

The mantle source for the Waiareka-Deborah volcanic field has not been investigated in any detail prior to our work. The data of Hoernle et al. (2006; Sr-Nd-Pb; n = 3) and Fulmer et al. (2010; Hf, n = 1), are here supplemented by Sr-Nd-Pb for 9 samples (Table 1). The new isotope data were obtained at the University of Cape Town in South Africa following the method reported by Scott et al. (2020).

The ⁸⁷Sr/⁸⁶Sr_{(m; measured)} values for sub-alkaline Waiareka-Deborah volcanic rocks cluster between 0.70321 and 0.70387, overlapping those for the Kakanui melanephelinite (0.70342). Low Rb/Sr means that age correction to 34 Ma makes little or no difference to the ⁸⁷Sr/⁸⁶Sr values (⁸⁷Sr/⁸⁶Sr_(i) sub-alkaline = 0.70310–0.70385; alkaline = 0.70339) (Figure 8A). The Tawhiroko and Moeraki sills contain locally abundant pyrometamorphosed Otago schist xenoliths in which Rb-rich micas appear to have been assimilated (Figure 4E), but the isotopic similarity between these samples and other sub-alkaline flows indicates that the assimilation process was localised and the isotopic compositions of the analysed samples were not significantly changed. Alternatively, lack of a geochemical signature indicating crustal contamination may indicate that schist-rich layers formed as late injections into an inflating sill (Marsh 1996).

Figure 8. Bulk rock isotopic data for the Waiareka-Deborah Volcanic Field compared to other volcanic fields in Otago and clinopyroxene in mantle xenoliths of those fields. A, ⁸⁷Sr/⁸⁶Sr_{(i; initial)} versus ¹⁴³Nd/¹⁴⁴Nd_(i) shows the Waiareka-Deborah Volcanic Field tends to be slightly more radiogenic in Sr than the bulk of the other intraplate provinces. B, ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁸Pb/²⁰⁴Pb_{(m; measured)} highlights the distinct nature of the Waiareka-Deborah Volcanic Field in ²⁰⁷Pb/²⁰⁴Pb isotopes to the other intraplate provinces and the mantle lithosphere, except for basalts in the Maniototo in the Dunedin Volcanic Group. C, ²⁰⁶Pb/²⁰⁴Pb versus ²⁰⁸Pb/²⁰⁴Pb_(m) illustrates the mostly less radiogenic nature of the mantle source to the Waiareka-Deborah Volcanic Group compared to other Otago intraplate provinces. Data sources are: Waiareka-Deborah Volcanic Field rocks (this study; Hoernle et al. 2006) and xenoliths (McCoy-West et al. 2016); Dunedin Volcanic Group rocks (Price et al. 2003; Hoernle et al. 2006; Sprung et al. 2007; Scanlan et al. 2020; Scott et al. 2020) and its peridotite xenoliths (Scott et al. 2014a; Scott et al. 2014b; McCoy-West et al. 2016; Dalton et al. 2017); Alpine Dike Swarm rocks (Barreiro and Cooper 1987) and peridotite xenoliths (Scott et al. 2014b, 2016); crustal Sr-Nd-Pb data are from Scanlan et al. (2018, 2020) and Scott et al. (2020).

The sub-alkaline ¹⁴³Nd/¹⁴⁴Nd_(m) data have a tight moderately radiogenic cluster (0.51284–0.51288; εNd = +4.7 to +3.9), with the melanephelinite plotting at the more radiogenic end of this dataset (0.51290; εNd = +5.1) (Figure 8A). Low Sm/Nd means that age correction to 34 Ma for all these rocks makes little difference (0.51281–0.51288; εNd = +4.7 to +3.3).

The Pb isotopes show little variation across the volcanic field but do indicate a subtle distinction between sub-alkaline rocks and the melanephelinite (Table 1). There are, however, insufficient high-quality U-Th-Pb trace element data to apply an age correction, so only uncorrected data are reported. ²⁰⁶Pb/²⁰⁴Pb of the sub-alkaline rocks (19.042–19.247) is less radiogenic than in the Kakanui melanephelinite (19.313), as are ²⁰⁸Pb/²⁰⁴Pb ratios (sub-alkaline = 38.763–38.921; melanephelinite = 38.948). The sub-alkaline ²⁰⁷Pb/²⁰⁴Pb data (15.617–15.649) overlap with those from melanephelinite (15.622), yet are distinct in ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb space (Figure 8B, C).

Clinopyroxene separates have been analysed for Sr, Nd and Pb from 8 peridotite xenoliths (Scott et al. 2014b, n = 7; McCoy-West et al. 2016, n = 1). This suite comprises two xenoliths from Round Hill, two from Alma and four from Kakanui. ⁸⁷Sr/⁸⁶Sr_(i) in the peridotite xenoliths (0.70242-0.70316) is less radiogenic than solidified host magmas, whereas ¹⁴³Nd/¹⁴⁴Nd_(i) data (0.51360–0.51285) overlap. The extremely unradiogenic Sr and radiogenic Nd of several samples testify to portions of this mantle lithosphere having been isotopically isolated for long periods of time, although most of the samples appear to preserve a metasomatic isotopic composition that has been shown to be widespread in the Otago mantle (Scott et al. 2014a; Scott et al. 2014b; McCoy-West et al. 2016; Dalton et al. 2017; Scott et al. 2020). Peridotite clinopyroxene Pb isotope data extend from extremely unradiogenic through to radiogenic values (²⁰⁶Pb/²⁰⁴Pb = 17.77–20.2, ²⁰⁷Pb/²⁰⁴Pb = 15.41–15.65 and ²⁰⁸Pb/²⁰⁴Pb = 37.47–39.46) and only the Kakanui melanephelinite overlaps with this array. In detail, however, the xenolith Pb array is pulled to unradiogenic composition by an extremely unradiogenic single sample; exclusion of this would mean that no volcanic component so far analysed in the Waiareka-Deborah Volcanic Field overlaps with the Pb isotopic composition of analysed North Otago mantle xenoliths.

Discussion

Shallow-marine volcanism and shallowly emplaced sills

The dominance of volcaniclastic deposits formed by largely shallow-submarine eruptions is atypical of most studied volcanic fields in intraplate continental settings. Volcanologically analogous submarine volcanic fields formed on the ocean floor, such as the Vestmann Islands in Iceland (Jakobsson 1968, 1979) or North Arch in Hawaii (e.g. Frey et al. 2000; Davis and Clague 2006), are difficult to investigate in detail. In terms of volcano morphology and the general setting of the Waiareka-Deborah volcanic field, the Vestmann Islands in Iceland (Jakobsson 1968, 1979) offer a partial analogue. The depth of the Vestmann Island shelf, ∼100-150 metres, is very similar to depths inferred for the Waiareka-Deborah (mostly ‘inner shelf’, but up to ∼200 m; Hicks 2014). In both fields, the shallowest depths would have been at the coastlines of temporary islands. Observations during the eruption of Surtsey, and of other surtseyan volcanoes, show quick growth to sea level and above to form islands, with rapid, often repeated, planation during or soon after eruption (Hoffmeister et al. 1929; Thorarinsson 1967; Johnson and Tuni 1987). After formation, the Waiareka-Deborah volcanoes would have been reduced to shoals at fair-weather wavebase or below, over the course of months or years.

A significant feature of this field is the extent and volume of shallowly emplaced sills. Individual sills have been mapped across tens of km, whereas the extent of significant deposits from single volcanic centres appears to only be a few km or less. Coombs et al. (1986) inferred that the ‘ … total volume of all the currently visible intrusives … may have reached 10 km³ .’ This is a volume that may be half that of known Waiareka-Deborah volcaniclastic deposits. The sills were emplaced into a sedimentary succession that was on the order of only ∼ 100 to perhaps several hundred metres thick (Thompson et al. 2014), which means the ∼ 50 m thickness of several sills is very large relative to the overlying host rocks. For example, the Tawhiroko sill at Moeraki is closely associated with deformed submarine pyroclastic deposits, quite possibly co-genetic, from which it is separated by tens of metres or less of marine mudstone. The thinness of the sills is in stark contrast to those in thick basinal deposits of the North Sea, for example (Hospers and Ediriweera 1991; Thomson and Schofield 2008); it is not clear whether this is simply the result of smaller volumes of magma delivered to the Waiareka-Deboorah sills, or reflects an additional control by the thin basinal host deposits. The thinness of the Oamaru region's sedimentary succession reflects a long period of slow deposition and occasional hiatuses during which hardgrounds formed (Thompson et al. 2014). Features that could have promoted sill emplacement include readily deformable near-surface deposits (Galland et al. 2018), or rheological differences among the sedimentary units (Kavanagh et al. 2006). Understanding why this thin sedimentary sequence captured such substantial proportions of the magma sent toward the surface during Waiareka-Deborah time is a topic for further research.

Mantle sources for Waiareka-Deborah magmatism

Together the new major element (Figure 6) and trace element (Figure 7) data emphasise that the majority of analysed Debora-Waiareka Volcanic Formation is distinct from the slightly younger Alpine Dike Swarm and Dunedin Volcanic Group, which are Cenozoic intraplate volcanic provinces in west and east Otago respectively (Figure 1 inset). The lower incompatible element abundances of the Waiareka-Deborah suite cannot easily be explained as a result of crustal contamination because, as has already been argued for New Zealand intraplate rocks (Sprung et al. 2007), the continental crust is enriched in many of these elements (e.g. Otago Schist composition in Figure 7) and assimilation should lead to magmas with higher incompatible element abundances. Furthermore, there is no clear mixing array between crust and mantle in Pb isotopic space (Figure 8C). While the Waiareka-Deborah data are similar to the cluster of transitional alkaline–sub-alkaline Maniototo basalts in the NW corner of Dunedin Volcanic Group (Figure 7), the Maniototo basalts lack positive Sr and P anomalies and have ubiquitous negative K anomalies (Figure 7A) (Scott et al. 2020). While the positive Sr anomalies in the Waiareka-Deborah Volcanic Field (Figure 7A) could be due to plagioclase accumulation, this seems unlikely because Eu, an element that is also preferentially taken up in feldspar, shows only a very weak positive anomaly (Figure 7B). There is also no correlation between Sr and Mg# (not shown), which means that the Sr and P anomalies are not tied to crystal fractionation. A weak positive correlation of Sr and P (not shown) could be due to apatite in the mantle source. The positive Sr and P, the high SiO₂ content, low incompatible elements and absence of ubiquitous K anomaly compared to other Otago primitive intraplate basalts indicates the Waiareka-Deborah Volcanic Field magmas may have different mantle sources.

A distinct mantle source to the Waiareka-Deborah Volcanic Field compared to other Otago intraplate provinces is supported by radiogenic isotope data. ⁸⁷Sr/⁸⁶Sr data show that Waiareka-Deborah rocks are more radiogenic than most of the Dunedin Volcanic Group and Alpine Dike Swarm (Figure 8). Although the higher ⁸⁷Sr/⁸⁶Sr of this province could be a result of seawater interaction during submarine emplacement, this could not account for the distinct Pb isotopes (Figure 8A, B) and in any case interaction with seawater should give more extreme Sr isotope ratios due to the elevated Eocene seawater Sr isotope composition (e.g. Nelson et al. 2004). If interaction with seawater controlled the isotopic composition, it would mean that the sills were contaminated despite not breaching the surface and would require interaction with marine pore water to an equivalent degree across the entire volcanic field. Another option is that the elevated Sr isotope ratios are due to Otago Schist assimilation by an alkaline magma; however, this would have required large amounts of Otago Schist (since it has a much lower Sr content than the volcanic rocks) (Scanlan et al. 2020) and, in any case, this process would not deplete the incompatible element abundances such as Rb, Ba, Th, U, K etc. (Figure 7). Furthermore, although the Waiareka-Deborah ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb data trend towards continental crust (Figure 8B), the ²⁰⁶Pb/²⁰⁴Pb versus ²⁰⁸Pb/²⁰⁴Pb data plot along the Northern Hemisphere Reference Line (Figure 8C). The mantle sources to this volcanic province are also distinct from almost all the fragments of peridotitic mantle that have been exhumed from beneath the West, East and North Otago, which represent snapshots of the composition of the lithospheric mantle. Therefore, unlike the magmas of the Dunedin Volcanic Group and Alpine Dike Swarm, which have Sr, Nd and Pb isotopic compositions overlapping the mantle lithosphere, the isotopic properties of the Waiareka-Deborah Volcanic Field magmas appear to require derivation from a lithospheric mantle component that is either very rare, or derivation from the asthenosphere and rapid ascent without interaction with the lithospheric mantle.

Peridotite geothermobarometry and comparison with geotherms indicates that the Otago New Zealand lithosphere reaches about ∼70 km depth (Scott et al. 2014a; Scott et al. 2014b). If this is the case, then partial melting in the asthenosphere should involve partitioning of elements into refractory garnet and this should be chemically detectable. Sprung et al. (2007) pointed out that Lu/Hf fractionates more than Sm/Nd during melting in the presence of garnet. The very small variation in Sm/Nd (0.22–0.27) but large variation in Lu/Hf (0.045–0.09) in the Waiareka-Deborah rocks is therefore consistent with garnet having participated during melting. Furthermore, the alkaline basalts in the Otago area have near-ubiquitous negative K and Ti anomalies that represent either derivation from a K, Ti-depleted mantle (e.g. Timm et al. 2010) or amphibole ± phlogopite were residual during melting in the mantle source (Panter et al. 2006; Sprung et al. 2007; Pilet et al. 2008); however, these hydrous phases are thermally stable at lithospheric mantle temperatures of less than ∼ 1200°C. The Waiareka-Deborah Volcanic Field sub-alkaline basalts lack ubiquitous K or Ti anomalies (Figure 7A) and this observation coupled with the distinct isotopes suggests that they lack the lithospheric mantle signature apparently prevalent in the other Otago intraplate provinces.

A further feature relevant to the petrogenesis of the Waiareka-Deborah Volcanic Field is that the alkaline Kakanui Mineral Breccia melanephelinite (1) plots in the same (or very similar in the case of Pb) isotopic space as the sub-alkaline rocks and is (2) distinct from almost all the other alkaline basalts in Otago (Figure 8). This observation, coupled with overlapping Ar-Ar age, implies that the alkaline magmatism has the same (or very similar) isotopic reservoir to the sub-alkaline components and indicates that the petrogenesis of both types is likely related. Very low degrees (several percent) partial melting of NZ metasomatised peridotite has been modelled to be capable of giving rise to the distinctive trace element composition of alkaline basalt (Scott et al. 2016). With increasing degrees of melting, the incompatible trace element abundances of successive magmas are lowered and this results in a trace element composition similar to that of a sub-alkaline basalt. However, while the Kakanui Mineral Breccia magma has a prominent positive P anomaly, it lacks a positive Sr anomaly and also has prominent depletions in K, Zr, Hf and Ti. These latter element properties, which typically are interpreted to reflect source characteristics of the magmas (e.g. Panter et al. 2006; Sprung et al. 2007), imply that it is unlikely that both types of magma could be simply derived from different degrees of melting of a single mantle source and requires that the alkaline magmas have a distinct component that was melted at the same time.

Although sub-alkaline magmas have been experimentally generated by melting of mantle pyroxenite (e.g. Pertermann and Hirschmann 2003), the high Cr (∼300 ppm) in Waiareka-Deborah volcanic rocks (excluding those of the differentiated sills) indicates that peridotite has played an important role in the melting process. Furthermore, although garnet pyroxenite must reside in the mantle beneath the Deborah-Waiareka Volcanic Field (Figure 5B), their Sr, Nd and Pb data are more similar to the metasomatised mantle peridotites and alkaline volcanic rocks elsewhere in Otago than sub-alkaline volcanic components (Scott, unpublished data). In any case, there is no pyroxenite melting model that would alone account for the formation of SiO₂-undersaturated alkaline magma (e.g. Pilet 2015). The Kakanui Mineral Breccia melanephelinite could represent a hybrid magma that originated from the asthenosphere but assimilated metasomatised lithospheric mantle (e.g. Sprung et al. 2007). However, if this were the case, then the high Sr content of the melanephelinite (1379 ppm) compared to the sub-alkaline rocks (average of 456 ppm) requires that the mantle lithosphere contaminant had (1) a similar ⁸⁷Sr/⁸⁶Sr composition to the sub-alkaline source because the much higher abundance Sr abundance would lead to a large isotopic contribution, and (2) had sufficient residual amphibole and/or phlogopite to generate the negative K and Ti anomalies. The petrogenesis of the alkaline components in this dominantly sub-alkaline volcanic province is unresolved.

Conclusions

The intraplate Waiareka-Deborah Volcanic Field surrounding Oamaru in the South Island preserves a superb record of Eocene-Oligocene submarine volcanism on the Zealandia continental shelf. The on-land extent is ∼890 km², and the offshore extent could, potentially, be nearly four times that area. The volcaniclastic beds, which are extensive and well-exposed along the coastline, show that the volcanic field formed from small volume, short duration eruptions, with volcanoes built from tephra deposited by eruption-fed density currents, and more widespread tephras emplaced from ash initially transported in subaerial eruption plumes (Figure 3). The magma also crystallised as sills, pillow lavas and dikes. The sills are extensive, reaching up to ∼25 km², were emplaced at shallow levels in the thin sedimentary veneer that covered Otago at the time of formation, and are commonly chemically differentiated. Volcaniclastic glass chemistry and bulk rock composition of sills and pillow lavas show that the volcanic field mainly erupted sub-alkaline magmas, although there were minor coeval alkaline components, the best known of which is the xenocryst- and xenolith-rich Kakanui melanephelinite. Radiogenic isotopic properties of the sills and dikes and one melanephelinite clast indicate that the mantle source for the magmas was distinct from the nearby intraplate alkaline Oligocene Alpine Dike Swarm and the adjacent and spatially overlapping Oligocene-Miocene Dunedin Volcanic Group. Intraplate volcanism in Otago therefore has multiple mantle sources.

Acknowledgments

Alan Bischoff, Alan Cooper, Alessio Pontesilli, Daphne Lee and Marco Brenna are thanked for comments on various drafts. Stephen Read helped with Figure 1. Jenni Hopkins and John Smellie are thanked for reviews.

Data availability statement

All data related to this paper are summarised in Table 1, or in the cited papers.

Disclosure statement

No potential conflict of interest was reported by the author(s).