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Journal of Petrology Volume 42 Number 9 Pages 1643-1683 2001
© Oxford University Press 2001
Petrology and Geochemistry of the Lamongan Volcanic Field, East Java, Indonesia: Primitive Sunda Arc Magmas in an Extensional Tectonic Setting?
DEPARTMENT OF EARTH SCIENCES, CAMBRIDGE UNIVERSITY, DOWNING STREET, CAMBRIDGE CB2 3EQ, UK
Received September 30, 1999; Revised typescript accepted February 26, 2001
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONNew geochemical data are presented from prehistoric and historical eruptive products of the Lamongan volcanic field (LVF), East Java; a region of the Sunda arc covering
260 km2 and containing 90 eruptive vents plus the historically active Lamongan volcano. LVF lavas include medium-K basalts and basaltic andesites from historical eruptions of Lamongan and prehistoric eruptions in the eastern LVF, along with a high-K suite represented by prehistoric deposits in the western LVF. Although lacking some of the characteristics of truly primary basalts, the least evolved lavas identified in the LVF have some of the lowest SiO2 contents (43 wt % SiO2) yet reported in Sunda arc volcanic rocks. Mass balance considerations indicate that two chemically distinct LVF magmas may be parental to suites currently being erupted from the neighbouring volcanoes, Semeru and Bromo. Lamongan’s historical lavas can be related to the medium-K andesitic products of Semeru by fractional crystallization, despite the former’s location at the same distance from the trench as Bromo, a high-K volcano. Extensional tectonics, possibly related to arc segmentation in the region of the LVF, creating conditions that promote the rapid ascent of parental magmas, is probably responsible for this and several other features of the complex. KEY WORDS: geochemistry; primitive magmas; Sunda arc; volcanic field; extensional tectonics
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INTRODUCTION |
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The generation and subsequent evolution of magma in subduction zone settings is widely acknowledged to be a multifarious process, involving possible inputs from the subducted oceanic lithosphere, subducted oceanic sediments and fluids, the asthenospheric and lithospheric portions of the mantle wedge above the subduction zone, and the arc crust (e.g. Tatsumi & Eggins, 1995
). Following melt generation, processes such as crystal fractionation, accumulation and crustal assimilation during transit to the surface may obscure the true nature of the magma source region. The effects of these processes on the compositions of erupted lavas must therefore be removed before models of arc magma genesis can be tested. For this reason it is desirable to sample magmas that have suffered minimal modification since segregation from their source region. The compositions of the primary or primitive magmas from different arcs provide a window on the diverse mechanisms of, and the components involved in, melt production at convergent margins, and as such are essential to our understanding of the subduction zone environment.
Many studies have highlighted the correlation between extensional tectonics in volcanic arcs and the eruption of lavas with primitive chemical characteristics (e.g. Luhr & Carmichael, 1981; Knittel & Oles, 1995; Bacon et al., 1997; Smith et al., 1997). In such settings, the development of thinned and fractured crust allows rising magma to erupt at the surface in a relatively unmodified state, rather than ponding at the crust–mantle boundary. The rapid passage to the surface of individual magma batches inhibits the formation of long-lived crustal magma chambers, and invariably leads to the formation of monogenetic vent fields comprising tens to thousands of individual eruptive centres (such as cinder cones and maars) in these environments [e.g. the Michoacán–Guanajuato Volcanic Field, Mexico (Luhr, 1997) and the Macolod Corridor, Philippines (Knittel & Oles, 1995)].
The Sunda arc in Indonesia is a mature island arc with several structural features indicative of an active extensional regime at certain points along its length, such as developing back-arc basins (e.g. the Madura Basin, East Java; Hamilton, 1979; Fig. 1). Furthermore, the present seismic strain field beneath the volcanic arc on Java is indicative of extension oriented at a high angle to the arc, a feature common to many convergent margins (Apperson, 1991). However, there have been few documented studies of monogenetic vent fields on the Sunda arc, which, as discussed above, are also symptomatic of intra-arc extension.
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This paper presents the results of a petrological and geochemical investigation of the Lamongan volcanic field (LVF) in East Java, which encompasses the historically active Lamongan volcano along with numerous cinder cones and maars (Carn, 2000). The surrounding region contains at least five recently active volcanoes (Fig. 1). The LVF sample suite covers most of the exposed vents and lava flows, and allows the characterization of the most primitive magmas in the complex. Modelling of fractional crystallization is used to place these lavas in a local and regional context. The data are also used to make preliminary judgements on the mechanisms of magma genesis in this part of the Sunda arc and on the possible effects of the regional tectonic regime on volcanism in the area.
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OVERVIEW OF LAMONGAN VOLCANO AND THE LVF |
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Tectonic and geological setting
Lamongan volcano (8·00°S, 113·34°E; altitude 1651 m) lies on the Sunda arc between the massifs of Tengger–Semeru and Iyang–Argapura, around 200 km above the Wadati–Benioff Zone (Fig. 1). Its location corresponds to one of the narrowest sections of the island of Java, and graben structures in the vicinity of the LVF suggest local extension (Carn, 1999
). The region has been little studied, and prior petrological studies are only cursory (Mulyana, 1989). The occurrence of three seismic crises in the area during the 20th century, each entailing the formation of east–west- and ENE–WSW-oriented tensional fractures on Lamongan’s western flanks (Matahelumual, 1990), led to the establishment of an observatory to monitor the volcano.
The Sunda arc extends from the Andaman Islands north of Sumatra to the island of Alor in the Banda Sea over a distance of >3000 km, and incorporates 76% of Indonesia’s Holocene volcanoes (Simkin & Siebert, 1994). Along much of the arc Indian Ocean crust is subducting beneath the Eurasian plate at a rate of 6 cm/yr (Hamilton, 1979), with the downgoing slab probably continuous from the surface to the lower mantle beneath Java (Puspito & Shimazaki, 1995; Widiyantoro & van der Hilst, 1996, 1997). This continuous slab is compatible with a regional subduction history that dates back to the Permian (Katili, 1975).
The Indian Ocean crust ages eastwards along the arc (Fig. 1), being around 130–135 Ma old south of Java (Widiyantoro & Hilst, 1996). In the eastern Sunda arc, the incursion of continental crust and crustal fragments on the Indo-Australian plate complicates the system (Fig. 1). East of Flores the progressive docking of Australian continental lithosphere onto the Banda arc since the Neogene (e.g. Rangin et al., 1990) has led to a cessation of volcanism at the arc front. This arc–continent collision is also believed to be responsible for incipient back-arc thrusting (from Bali eastwards; Fig. 1) and extensional features on the arc (e.g. Silver et al., 1983; Charlton, 1991).
The Sunda arc has existed in roughly its current configuration for around 5 m.y., although an active volcanic arc has probably persisted in the region since at least the Eocene (Katili, 1975; Hamilton, 1979; Rangin et al., 1990). The Quaternary volcanoes generally overlie Late Miocene–Pliocene volcanic rocks and sediments (Soeria-Atmadja et al., 1994). The changing nature of the arc crust, from continental in Sumatra to transitional in Java and oceanic in Flores, gives rise to along-arc variation in the style and composition of volcanism (Whitford et al., 1979). Silicic volcanism and large calderas (e.g. Lake Toba) on Sumatra give way to mafic–intermediate volcanism on Java, where the crust is of near-continental thickness but comprises mélanges and mafic–intermediate plutons. From Bali to Sumbawa (Fig. 1) a mature oceanic island arc has developed (Hamilton, 1988). Across-arc geochemical variations, primarily in K and 87Sr/86Sr, have also been documented (Whitford & Nicholls, 1976; Whitford et al., 1979).
Lamongan volcano and the LVF
Lamongan volcano is the only historically active vent in the LVF, a region covering some 260 km2 occupied by numerous prehistoric eruptive centres including around 61 cinder cones and 29 maars (Fig. 2). Two prehistoric vents, Gunung Tarub and Gunung Tjupu, complete the central edifice (Lamongan–Tarub–Tjupu or LTT) of which Lamongan forms the southwestern segment (Figs 2 and 3). Carn (2000) presented an analysis of the physical aspects of the LVF, and his compilation of volumetric data for the cinder cones, maars and lava flows of the complex is given in Table 1. The number of vents in the LVF is exceptional for an Indonesian volcano, exceeding the 35 scoria cones reported from Gunung Slamet in central Java (Vukadinovic & Sutawidjaja, 1995), although it is somewhat fewer than that observed in larger cinder cone fields elsewhere (e.g. Mexico; Luhr, 1997). Vent distributions in the LVF suggest a strong regional tectonic influence on prehistoric volcanism in the field, which may have involved fissure-style eruptions (Carn, 1999, 2000).
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Lamongan’s first recorded historical activity occurred in 1799. This heralded 100 years of unrest during which the volcano erupted up to 15 lava flows, the last in 1898 (Matahelumual, 1990; Said, 1992; Simkin & Siebert, 1994). The steady-state eruptive rate at Lamongan from 1843 to 1898 was 0·03 m3/s, and the age of the volcano has been estimated at 13–40 ka, assuming a constant time-averaged eruptive rate over its lifetime (Carn, 2000). With the exception of seismic crises on the volcano’s flanks in 1925, 1978, 1985 and 1988–1989 (Matahelumual, 1990), Lamongan has been dormant since 1898.
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FIELD OBSERVATIONS AND SAMPLING |
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The LVF is situated
8° south of the equator and much of the field is well vegetated. All the observed lava flows in the LVF (historical and prehistoric) are of a’a, and are several metres thick and up to 2–3 km long (Fig. 3). Cross-sections of individual prehistoric flows, 10–30 m thick, were found in outlying stream sections and in the vertical walls of maar craters although they often proved difficult to trace beyond these isolated exposures. Several of the dry maars were observed to contain small cinder cones (Fig. 2) indicating a sustained eruption of magma following maar formation. The sample suite collected from the LVF is representative of most of the eruptive units and monogenetic vents in the complex (Figs 2 and 3). Samples were categorized as: (1) historical lavas (erupted from Lamongan volcano, 1799–1898); (2) cinder or spatter cones [including lava flows clearly associated with cones (mapped in Fig. 2) and cones within maars]; (3) prehistoric lava flows (flows not covered by the two preceding categories); (4) maar deposits (Table 2). No particular genetic relationships between samples in each group are implied by this classification. Samples were further subdivided on the basis of vent morphology and location, as detailed in a later section. The majority of LVF lavas are basalts and basaltic andesites, with subordinate picrobasalts, basanites, trachybasalts and trachyandesites (Table 2).
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Samples from cinder or spatter cones were invariably scoriae or spindle bombs found close to the surface. Dense vegetation in the SE of the region precluded detailed exploration, but there is little evidence for young vents in this area (Bronto et al., 1986). Maar deposits were poorly preserved and juvenile material was difficult to identify. The provenance of maar samples is hence largely unknown; some material may have been reworked by the disruption of subsurface lava flows. For comparison, several new analyses from volcanoes adjacent to the LVF (Semeru and Bromo) are also presented (Fig. 1; Table 2).
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ANALYTICAL TECHNIQUES |
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Whole-rock samples were analysed for major and trace elements by X-ray fluorescence (XRF) at the Open University, Milton Keynes, on an ARL 8420+ wavelength-dispersive XRF spectrometer. Fresh cores of lavas were washed in deionized water and dried before crushing in an agate grinder. Pressed powder pellets for trace element determination were prepared in Cambridge, and fused glass discs for major element analyses were prepared in Milton Keynes. Data quality was monitored by running in-house igneous rock standards with each sample batch. Trace element detection limits are given in Table 2. Mineral analyses were performed on carbon-coated polished slides using a Cameca SX-50 electron microprobe in Cambridge. Analytical conditions were a 10 kV accelerating voltage, beam currents of 4–15 nA and a beam diameter of 5 µm; all analyses were energy dispersive. Precision, monitored using various standards, is generally within 1–2% for major elements. To determine Sr isotopic ratios,
150 mg aliquots of rock powder were dissolved in HF–HNO3 and the resulting solution was evaporated until dry. After further dissolution in HNO3 the solution was evaporated again; the samples were then redissolved in HCl. After centrifuging, samples were loaded onto cation-exchange resin and Sr was eluted using HCl. The dried residue was loaded in HCl onto an outgassed Ta filament and analysed on a VG Sector 54 mass spectrometer in Cambridge. Ratios were corrected for mass fractionation by normalizing to 86Sr/88Sr = 0·1194. Repeat analyses of NBS 987 gave a mean value of 0·710263. |
PETROGRAPHY |
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Historical lavas
Historical lavas from Lamongan volcano typically contain phenocrysts of plagioclase (Plag), clinopyroxene (Cpx), olivine (Ol) and Ti-magnetite (Mag) in a crystal-rich, seriate-textured groundmass with similar mineralogy and variable amounts of glass (often devitrified). The petrography of the lavas is generally rather uniform throughout the series. Modal analyses (point-counted) for 10 historical lavas are listed in Table 3.
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Subhedral Plag is the predominant phase in all samples, accounting for up to 70% of the modal mineralogy (Table 3). It typically occurs as complex aggregates or glomerocrysts (up to 3–4 mm square; associated with Cpx and Mag), and often exhibits fine-scale concentric zoning, sieve texture and resorption features (e.g. embayed rims), and zoned inclusions suggesting a complex history. Microphenocrysts often show similar features. Inclusions of pyroxene and Mag are often observed in Plag. Groundmass alignments of Plag are rarely observed and are never well developed. Phenocrysts range up to 3–4 mm in size.
Cpx also occurs in aggregates (with Plag, Fe–Ti oxides and Ol), although less frequently than Plag. It is typically subhedral, concentrically zoned, and occasionally displays twin lamellae; phenocrysts are up to 2 mm in size. Embayed and sieve-textured Cpx (and rare skeletal grains) is also present in some samples. Inclusions of Plag are rare, but zones of Fe–Ti oxide inclusions are frequent; a common texture observed in the historical lavas (and in other samples; see below) is the mantling of Cpx phenocrysts by small Fe–Ti oxide crystals. Low-Ca pyroxene is present only as a minor phase in the more evolved lavas (e.g. LAM1561). It occurs as resorbed crystal cores or as rims enveloping Cpx phenocrysts.
Ol is typically less abundant than Plag and Cpx in the historical lavas (Table 3). It occurs as phenocrysts (up to 1 mm across) and microphenocrysts (generally <0·5 mm across) that are frequently embayed, and it is also present in the groundmass of several samples.
Also observed in several samples (e.g. LAM18981, LAM18831, LAM1070) are clots or nodules containing Plag, Cpx, Mag and occasional Ol, up to 0·8 mm in size. The pyroxene in these clots is generally unzoned and often shows a subophitic or ophitic texture, distinguishing them from the ubiquitous glomerocrysts that have similar constituents. Several nodules show disequilibrium features such as sieve-texture and resorbed boundaries. These nodules are interpreted as cumulate material derived from the edges of an evolving magma chamber.
Plag glomerocrysts are less common in the younger lavas (LAM18982, LAM18832; Table 3), suggesting that a flotation mechanism (e.g. Campbell et al., 1978) may have controlled their distribution in the chamber and biased their occurrence to earlier-erupted flows.
Cinder or spatter cones
Samples from the LVF cinder or spatter cones display a wide range of petrographic features. Following Carn (2000), the cones are subdivided morphologically into a young series (centuries old) and an old series (millennia old; Fig. 2). The former is referred to as the kerucut volkanik muda (KVM) series and the latter as the kerucut volkanik tua (KVT) series. Cinder cones observed within maars belong to the KVM series. The KVT cones are generally situated in the western part of the LVF (Fig. 2). The modal mineralogy of three KVM samples is given in Table 3.
The KVM lavas contain phenocrysts of Plag, Cpx, Fe–Ti oxides and Ol. Very rare crystals of oxy-hornblende are also present in two scoria samples (ALA1161 and GCK2961). Many of the KVM cones are spatter ramparts associated with flank lava flows (Fig. 2), and samples from these flows (e.g. GCK2960, GKL256) are similar in appearance and modal mineralogy to the historical lavas (Table 3). Large Cpx phenocrysts (up to 7 mm long, often euhedral) are common in the KVM lava flows, typically with inclusions of Fe–Ti oxides, often concentrated at crystal rims (as observed in the historical lavas). Some Cpx shows resorption features (e.g. embayments, eroded rims, sieve texture) and exsolution lamellae are common. Ol phenocrysts (up to 6 mm long) and microphenocrysts (<1·5 mm long) are usually embayed and occasionally skeletal. Plag is abundant and sometimes occurs as glomerocrysts, but less commonly than observed in the historical samples (Table 3), and disequilibrium textures in Plag are also present, but less widespread in the KVM samples. Plag phenocrysts are typically a maximum of 1·5 mm in length. Fe–Ti oxides are abundant, particularly in the groundmass; phenocrysts are up to 2 mm in size. Cumulate clots of Plag, Cpx and Fe–Ti oxides, similar to those discerned in the historical lavas, are also evident in the KVM lava flow samples. There is a general trend of decreasing phenocryst dimensions with distance from the central (LTT) edifice.
Other KVM samples are basaltic scoriae or small bombs, and hence contain more groundmass glass than the associated lava flows (e.g. ALA1161, Table 3). The scoriae have the same mineralogy as the lavas (with the exception of the two samples that contain minor oxy-hornblende), although phenocrysts are generally smaller in the former; typically <1 mm with rare Cpx up to 5 mm and Plag up to 3–4 mm in length. Other textural features are similar, such as the mantling of Cpx phenocrysts by Fe–Ti oxide inclusions, the presence of cumulate nodules (maximum 5 mm diameter; largely Cpx and Fe–Ti oxides), and some resorptional features in Ol, Cpx and Plag crystals. Glomerocrysts are typically absent or rare and disequilibrium textures are less common than in the historical samples.
KVT samples are petrographically diverse but some spatial patterns can be discerned. Cones in the western LVF, west of the town of Klakah (Fig. 2), are particularly distinct and represent some of the oldest vents in the complex (Carn, 2000). Samples from these cones are generally glassy and contain Plag, Ol and minor Cpx in order of decreasing abundance. Plag typically exists as microlites (normally aligned) or microphenocrysts, often zoned and occasionally showing disequilibrium textures in larger crystals, which may have fresh mantles. Glomerocrysts of Plag up to 3–4 mm in diameter are also observed in some samples, and individual phenocrysts range up to 3 mm in length. Ol occurs as microphenocrysts (up to 1·5 mm in size), which are frequently embayed or skeletal. Cpx appears rare in these samples and is generally restricted to the groundmass.
Other KVT samples form localized groups with distinctive petrography, although they are generally akin to the KVM lavas. Textural similarities are observed between samples from collinear vents (e.g. GMI15072 and GKN3060; both highly Plag-phyric).
Prehistoric lavas
Prehistoric lava flows (PLFs) are distributed throughout the LVF although they are rarely well exposed. Samples from outlying flows (PLF1; e.g. TAR1361, KAP2062, RBI1162; Fig. 2) are typically Plag-phyric with a glassy matrix. Cpx and Ol are present as microphenocrysts (1 mm) although Ol is often unstable. Plag crystals are usually flow-aligned and show rare disequilibrium textures. Some clots (1 mm diameter) of Cpx, Fe–Ti oxides and Plag are also present, and the groundmass is often rich in Fe–Ti oxides.
The other PLF samples form two distinct groups. One set (PLF2) crops out around much of the flank region of the LTT edifice, usually in the walls of maar craters as thick flows (e.g. RWG116, RJN278, RAG1962, RLG1262; Figs 2 and 3); these lavas are highly porphyritic, containing phenocrysts of Ol, Plag and Cpx in a coarse groundmass (same mineralogy, plus Fe–Ti oxides). Ol is typically anhedral and unstable (skeletal in places), occurring as microphenocrysts and phenocrysts up to 3 mm in length and occasionally in nodules (with variable amounts of Fe–Ti oxides, Cpx and Plag) several millimetres in diameter. Plag often occurs as glomerocrysts (3–4 mm in size), and frequently displays sieve-texture and fine-scale compositional zoning, the latter sometimes mantling the former. Cpx phenocrysts (3–6 mm) and glomerocrysts (up to 1 cm long) are less abundant, sometimes euhedral, but generally unstable. Some contain inclusions of Ol, Plag and Fe–Ti oxides (concentrated at the edges of the host crystal).
The other group (PLF3) crops out at a slightly higher stratigraphic level in flows with seldom pristine morphologies, largely on the flanks of Lamongan (e.g. LAO306, LAO1207, TAR2760; Figs 2 and 3). Their petrography exhibits many similarities with the historical samples from Lamongan.
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MINERALOGY |
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The major phenocryst phases in LVF lavas are Ol (Fig. 4), Cpx (Fig. 4), Plag (Fig. 5) and Fe–Ti oxides. Representative mineral analyses are presented in Tables 4–6, and a summary of observed mineral compositions is given in Table 7. The majority of Fe–Ti oxide crystals analysed fall on the magnetite–ulvöspinel solid solution (titanomagnetite), with TiO2 contents as high as 17 wt % but typically around 7–10 wt %.
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Olivine
Forsterite (Fo) contents of Ol phenocrysts and microphenocrysts in LVF lavas range from Fo29 to Fo79, with considerable variation often observed on the scale of individual samples (Table 7). This variation is generally greater in the historical lavas. Many crystals display normal zoning (Table 4). The more Fo-rich crystals are generally found in cinder cone samples, and Fo contents are typically highest in glomerocrysts and crystals found in cumulate clots or nodules, with microphenocrysts typically more iron rich. Disequilibrium textures in Ol are seen in a range of compositions (Table 4). CaO contents rarely exceed 0·3 wt %, which is a typical value for basaltic olivines (Jurewicz & Watson, 1988) and higher than that found in mantle olivines (<0·1 wt %; e.g. Foden, 1983).
There is widespread evidence for Ol accumulation in LVF lavas. Ol phenocrysts and microphenocrysts are typically not in equilibrium with their host rock, with Fo contents generally being too low for their respective bulk-rock mg-number; this is particularly evident in the historical lavas and PLFs (Fig. 6). This may be due to elevation of the measured Fe2O3/FeO (wt %) ratio by post-eruption alteration and oxidation, which would result in artificially high mg-numbers. However, most lavas were not visibly altered, with the exception of some PLF2 samples that show iddingsite rims on Ol phenocrysts (e.g. RAG1962), which may explain the more extensive disequilibrium in these samples (Fig. 6). Assuming an Fe2O3/FeO ratio of 0·2 still places the majority of historical samples and PLFs in disequilibrium, although the cinder cones (largely KVM) plot close to the high-pressure equilibrium field (Fig. 6). Cinder cone Ol also shows the narrowest range of Fo contents, particularly sample ALA1161 (KVM), whereas the PLF2 samples (RWG116, RAG1962; Fig. 6) show the clearest evidence for Ol accumulation. Mean Fo contents in Ol are 63% (historical lavas), 74% (cinder cones) and 62% (PLFs).
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), cinder cones (
, 
) and PLFs (
, 
); symbols represent phenocrysts and microphenocrysts, respectively. Historical lavas are labelled with eruption dates (where known); cinder cones and PLFs are labelled with sample names. Dotted lines show where the data would plot if mg-numbers were recalculated assuming Fe2O3/FeO = 0·2. The shallow-pressure (SP) and high-pressure (HP) equilibrium fields for basaltic magma (Roeder & Emslie, 1970
Pyroxene
The vast majority of pyroxenes are typical augites (Fig. 4), with minor low-Ca pyroxene observed only in the more evolved historical lavas, and rare ferroaugite. Where present, low-Ca pyroxene is occasionally mantled by Cpx and often unstable. As mentioned above, inclusions of Mag are very common towards the rims of pyroxene phenocrysts. This may be a reaction texture resulting from the breakdown of unstable Fe-rich pyroxene, or an oxidation effect. Pyroxenes can contain high concentrations of structurally bound OH (up to 110 ppm H2O), particularly mantle pyroxenes (e.g. Skogby, 1994). Such OH may provide fuel for oxidation–dehydroxylation reactions that occur during the formation of oxy-hornblende and convert Fe2+ to Fe3+. Phenocryst rims and groundmass crystals are typically more Fe rich. CaO contents in excess of 23 wt % (salite) are not uncommon in samples from cinder cones and PLFs (Table 5). Chromian pyroxene is absent, with Cr2O3 typically <0·1 wt %.
Pyroxenes from cinder cone and PLF samples are distinguished by their high Al contents, often in excess of 5 wt %, and also by high TiO2 (1 wt % or more; Table 5). Al2O3 concentrations in clinopyroxene from subcontinental and suboceanic peridotites range from 2 to 8 wt %, with decreasing Al associated with increasing proportions of coexisting melt (Seyler & Bonatti, 1994). Similar augites and Ti-rich aluminous salites to those found in the LVF have been reported in alkaline (potassic) basalts from Ringgit–Beser, an extinct volcanic complex situated on the north coast of East Java (Fig. 1; Edwards et al., 1994), and in calc-alkaline basalts from behind the volcanic front along the Sangihe arc north of Sulawesi (Morrice & Gill, 1986).
Plagioclase
Plag exhibits a wide compositional range in many samples, with a maximum anorthite (An) content of 96 mol % (Table 7; Fig. 5). The more An-rich compositions are found in glomerocrysts and as phenocryst cores (Table 6), and are more common in cinder cone samples.
Highly calcic plagioclase is common in high-Al arc basalts, with An content dependent upon the H2O content, crystallization pressure, Al2O3 content and CaO/Na2O value of the melt (e.g. Panjasawatwong et al., 1995). Melt CaO/Na2O exerts the strongest control on plagioclase An content, with values of CaO/Na2O > 8 required to produce highly calcic (An95) plagioclase as observed in the LVF samples. The highest CaO/Na2O ratios in the LVF suite (6–7) occur in the KVM lavas from the northern and western flanks of LTT (Fig. 2), which also contain An-rich Plag (Table 7), but these are probably not true liquid compositions. High values (CaO/Na2O 7) also occur in the PLF2 samples, although these are highly phyric.
Disequilibrium textures (embayed crystals, sieve texture, etc.) are common in Plag (particularly in the historical lavas and the PLF2–PLF3 samples) but are not restricted to a particular compositional range. Zoning of several types (normal, reverse, oscillatory) is prevalent in Plag, especially in the historical lavas (Table 7). Sieve-textured or cellular Plag can form through resorption (e.g. as a result of decreasing P, increasing PH2O, or mixing with hotter magma) or rapid crystal growth under vapour-saturated conditions or after mixing with cooler magma (e.g. Tsuchiyama, 1985; Morrice & Gill, 1986). These processes can be distinguished using textural evidence, because resorption leads to cellular zones with diffuse inner boundaries, whereas rapid growth results in zones with sharp borders, often associated with oscillatory zoning. The latter appear to dominate in the LVF lavas, although the cellular cores of Plag are often more diffuse.
Amphibole
The occurrence of amphibole in the LVF lavas is restricted to rare phenocrysts of oxy-hornblende in two KVM samples. These crystals were not analysed by electron microprobe. They have narrow rims of fine-grained oxides, with dark brown pleochroic interiors. The relatively high oxide content of some samples may indicate widespread oxidation and breakdown of amphibole in LVF magmas, which is only rarely preserved.
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GEOCHEMISTRY |
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Overview
Whole-rock major and trace element XRF data for the LVF suite are reported in Table 2. In Fig. 7, the LVF data are compared with existing analyses from other active and extinct Indonesian volcanoes. The least evolved LVF basalts (
43 wt % SiO2) represent some of the most SiO2-poor lavas reported from the archipelago to date, considerably lower than the mean SiO2 of 55·5 wt %. Although many of the LVF lavas have higher MgO contents than the relatively low Indonesian average of 4·5 wt % (calculated from the data in Fig. 7b), they also have rather low MgO at a given value of SiO2 in comparison with other Indonesian volcanoes. K2O contents in LVF samples are lower than the mean of 2·5 wt % for the data in Fig. 7a.
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Variations in K2O content broadly divide the LVF suite into two series, a medium-K series and a high-K calc-alkaline series (Fig. 8a). In Fig. 8, the LVF lavas are plotted along with data from Semeru and Bromo (Fig. 1) and from Iliboleng and Lewotolo, which lie east of Flores and north of the arc–continent collision zone (Fig. 1). The bulk of the LVF samples, including the historical lavas, most of the KVM series and the PLF2 and PLF3 samples, fall in the medium-K field and form an array that continues to higher SiO2 through the Semeru lavas. Crystal accumulation in the PLF2 series may have displaced them towards transitional tholeiitic compositions (Fig. 8a). The KVT series, PLF1 samples and the maar deposits generally lie in the high-K field, along with the rocks from the Bromo–Tengger complex and the eastern Sunda arc (Iliboleng, Lewotolo). Similar patterns emerge from a total alkali–silica plot (Fig. 8b), wherein the majority of LVF lavas are classified as typical basalts with subordinate picrobasalts (KVM series), basaltic andesites (including the oldest historical flows), trachybasalts and basaltic trachyandesites.
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)], Semeru (
), Bromo–Tengger (
As is commonly observed in subduction-related volcanic rock classification (e.g. Vukadinovic & Sutawidjaja, 1995), the use of different parameters to categorize the LVF lavas gives slightly contradictory results. On an AFM diagram (Fig. 9) the LVF samples straddle the tholeiitic–calc-alkaline boundary with a definite tholeiitic tendency, showing marked Fe enrichment with respect to the Semeru and Bromo–Tengger lavas, particularly at intermediate compositions. This is contrary to the K2O plot (Fig. 8a), but as Vukadinovic & Sutawidjaja (1995) pointed out, the two methods of discrimination probably relate to mutually exclusive aspects of the lava suite, namely its source (K2O) and its crystallization history (AFM). The LVF samples are thus relatively and variably K rich (probably inherited from the source) and also show Fe enrichment (probably related to magma chamber processes).
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Major and trace element variations
Harker diagrams of major element variations with SiO2 in the LVF suite are shown in Fig. 10. Both MgO and CaO behave compatibly throughout the entire series, indicating the influence of Ol and/or Cpx removal. A group of three samples with 10–11 wt % MgO (PLF2 and maar deposits) have accumulated significant Ol; the other lavas have MgO <7·5 wt %. The KVM series show the highest MgO values, and the early historical flows (e.g. the 1821 flow) and more evolved KVT rocks the lowest (2·7 wt %). CaO also peaks in the most mafic KVM samples at 13·4 wt %.
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Levels of Fe2O3* and TiO2 show similar patterns to MgO and CaO until 51 wt % SiO2 when the data become more scattered, possibly reflecting the involvement of two distinct magma series (most clearly defined in the TiO2 diagram). Elevated Fe2O3* (15–16 wt %) and TiO2 (up to 2 wt %) contents are seen in the PLF1 and some KVT samples, and in several maar deposits; all these samples also show relatively high K2O and P2O5 (Fig. 10). The latter are both incompatible in the LVF suite and there is no evidence for appreciable apatite fractionation, although there is a slight decrease in P2O5 with increasing SiO2 in the high-K group.
The LVF lavas are highly aluminous (14·5–20 wt % Al2O3). There is considerable scatter in Al2O3 contents, probably as a result of variations in Plag abundance, as the most Plag-rich sample is also the most aluminous (GKN3060; 20·2 wt % Al2O3). Despite this, Al2O3 appears to increase initially in the LVF lavas, until an inflexion at 50–51 wt % SiO2 after which there is a general decline. This may also reflect multiple magma series, or possibly the onset of significant Plag fractionation.
Variations in trace element concentrations with SiO2 are shown in Fig. 11. Rb contents show similar patterns to K2O, with the historical lavas, PLF2, PLF3 and most of the KVM samples forming a linear array with lower Rb (6–18 ppm), and a more diffuse group with higher Rb (21–77 ppm). The latter group comprises a small number of KVM and KVT samples at low SiO2 and relatively high Rb that are closely related in space (samples with prefixes GCK, GKG, GMA and GKN, situated NE of Klakah; Fig. 2), and a cluster at higher SiO2 (>50 wt %) consisting of the PLF1 series, the KVT lavas from the western LVF and many of the maar deposits (Fig. 11).
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The relatively enriched, more evolved group of samples is also evident in the Ba, Zr, Nb, Y and (to a lesser extent) Sr plots (Fig. 11). Sr contents (300–600 ppm) show considerable scatter, as for Al2O3, suggesting the influence of Plag accumulation (e.g. Vukadinovic, 1993). The ‘low-SiO2, high-Rb’ samples mentioned above also have slightly elevated Sr, suggesting that these magmas may have suffered similar contamination.
Ba contents show a wide range (200–1000 ppm) and also define a linear array with relatively low concentrations increasing with SiO2, and a more enriched cluster at >50 wt % SiO2. Concentrations in several PLF samples displaced below the linear array (Fig. 11) may have been diluted through the accumulation of ferromagnesian phenocrysts. Similar patterns are displayed by Zr, Y and Nb, with the divergent trend or enriched group at >50 wt % SiO2 particularly evident. Zr contents are rather depleted in the low-SiO2 samples (<50 ppm) but increase to 75 ppm at 54–55 wt % SiO2 in the most evolved historical lavas, and attain up to 226 ppm in some PLF1 and KVT rocks. Y varies from 15 to 27 ppm with increasing SiO2 in the linear array and up to 58 ppm in the enriched group. Niobium contents are also relatively low in the LVF lavas, with many samples close to detection limits (Table 2) and the highest contents (10–11 ppm) recorded in a KVT sample and maar deposit. In the enriched group of samples, Zr, Y, Nb and possibly Rb and Sr all appear to decrease with increasing SiO2 (Fig. 11).
Both Sc (15–60 ppm) and V (100–670 ppm) contents exhibit a steady decline with increasing SiO2, although there is also evidence for an enriched group of samples at >50 wt % SiO2, as observed in the incompatible element plots (Fig. 11). This compatibility of Sc and V in the LVF lavas attests to the influence of Cpx removal throughout the suite, with other deviations (particularly in Sc) probably attributable to varying modal proportions of the mineral.
Ni (<60 ppm) and Cr (<160 ppm) contents are both low in the LVF lavas, peaking in PLF2 samples that have visibly accumulated Ol (Fig. 11). Low Cr values imply minimal contamination by chromian spinel (e.g. in Ol) or by Cr-rich pyroxene. The group of KVT and PLF1 basalts that are relatively enriched in incompatible elements also have slightly elevated Ni (30–40 ppm) but the values are still low. The historical lavas and some KVM samples are particularly depleted in Ni and Cr, even where Sc and V are elevated, which suggests that even the least evolved lavas have undergone some Ol fractionation. This is surprising given their low SiO2 contents, and another possibility is that low Ni and Cr were inherited from the source.
Figure 12 shows N-MORB normalized trace element abundances in LVF lavas. All of the samples show a prominent trough at Nb, which is characteristic of subduction zone magmas. This depletion of Nb relative to the large-ion lithophile elements (LILE; e.g. Rb, Ba, K) is attributed primarily to two processes: (1) the addition of an LILE-enriched, Nb-poor fluid component to the mantle wedge; (2) the preferential retention of Nb in amphibole relative to other phases present in the mantle source (e.g. Borg et al., 1997). Similar processes are inferred to explain the general depletion of the high field strength elements (HFSE) Zr, Ti and Y with respect to LILE in arc magmas (e.g. Pearce & Peate, 1995).
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Inter-element trends for the LVF samples generally resemble a typical primary island-arc basalt (IAB) profile, but with more enriched values of most elements, although this is of course dependent on the IAB values used. Historical samples form the most homogeneous group, with notable enrichments in Rb, Ba, K, Pb, P, Ti and Y compared with IAB, and levels of Ba and Pb narrowly exceeding OIB (Fig. 12a). Ba enrichments in Javanese volcanic rocks are probably related to the input of Ba-rich oceanic sediments into the subduction zone (e.g. Plank & Langmuir, 1993).
The cinder cone lavas display varying levels of enrichment, although some of the variation will be due to crystal fractionation (Fig. 12b). KVT samples from the western LVF are typically most enriched, up to 10x primary IAB for Rb, and with relatively high values (compared with IAB) for all other elements except Sr (indicating Plag fractionation), exceeding or approaching ocean-island basalt (OIB) values in most cases. The KVM rocks and the cinder cones NE of Klakah (Fig. 2) generally represent the least enriched (i.e. most IAB-like) samples in the LVF. The latter have an interesting profile in Fig. 12b, showing significant Rb enrichment relative to IAB without a corresponding Ba anomaly. They are also depleted in Nb, P, Zr and, to a lesser extent, Sr, although enriched in Ti and Y, compared with IAB.
The PLF data display a similar range to the cinder cone samples, although there are two discrete groups rather than a continuum of compositions (Fig. 12c). The upper group is highly enriched in Rb, Ba, K, Pb, P, Zr, Ti and Y with respect to typical IAB and consists of the PLF1 series; these samples also show a small negative Sr anomaly similar to the KVT cones in the western LVF. The less enriched group comprises the PLF2 and PLF3 series, with the former exhibiting slight depletions in Nb, P and Zr, and enrichments in Ti and Y, relative to IAB (Fig. 12c). The only element that appears to be relatively constant throughout the suite is Sr, which is typically at or close to typical IAB abundance.
In summary, all of the LVF samples are enriched in Rb, Ba, K, Pb, Ti and Y to some degree relative to a typical primary IAB. Most of the lavas carry the typical subduction signature of Nb and HFSE depletion and LILE enrichment, with the cinder cone and PLF samples showing the greatest variability. Several of the KVM and PLF2 samples show low abundances of Nb, P and Zr relative to typical primary IAB and N-MORB, but without corresponding depletions in Ti and Y, and some cones appear to preserve localized geochemical heterogeneities such as enrichment in Rb relative to Ba. Abundances of Rb, P and Zr show the greatest, and Sr the smallest, range. A similar consistency of Sr abundance is seen in lavas from Gunung Slamet in central Java (e.g. Vukadinovic & Sutawidjaja, 1995).
Sr isotopes
The limited number of Sr isotope ratios available occupy a fairly narrow range (0·7042–0·70445; Table 2), with the lowest values close to the Sunda arc minimum (e.g. Whitford, 1975). This is consistent with the general decline in 87Sr/86Sr from West Java to Bali (Fig. 1).
The 87Sr/86Sr ratios do not correlate with SiO2 or Rb/Sr, although the highest ratio occurs in a KVM sample with anomalously high Rb and relatively high Rb/Sr (GCK2961; Table 2), which suggests source enrichment in this region. The lowest ratios occur in the historical lavas.
Principal component analysis
To clarify the complex patterns presented on binary diagrams by the large number of samples and elements analysed, the major and trace element XRF data from the LVF suite have been subjected to principal component analysis (PCA). PCA locates the directions of maximum variance in a multivariate dataset by finding the eigen values and eigenvectors of the data correlation matrix, and selects the combination of the initial variables responsible for this variance (e.g. Le Maitre, 1968; Albarède, 1995). In this analysis the historical lavas, cinder cones and PLFs have been treated as separate groups.
The first three principal components (PCs) of the historical lava data, which represent 85% of the total variance, are depicted in Fig. 13. The horizontal spread of these data (PC1) is generally related to the eruption age of the lavas, with PC2 probably picking out subtle variations in phenocryst content.
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At opposite extremes of the PC1 axis, we find the oldest (1821) and youngest (1898) historical samples. The former correlate with the incompatible elements (Rb, Zr, Ba, etc.) and relatively evolved components (SiO2, K2O, Na2O; e.g. alkali feldspar), whereas the latter correlate with a ‘pyroxene and Fe–Ti oxide’ component (i.e. MgO, Fe2O3, TiO2, CaO, V, Sc; Fig. 13). This probably implies that the magma chamber feeding the 19th-century eruptions at Lamongan was stratified, with the relatively evolved basalts at the top of the chamber erupted first (1821) and the more mafic magmas at the chamber base, enriched in dense ferromagnesian phenocrysts, erupted last (1898) and at the lowest altitude (Fig. 3).
Lavas erupted in 1864 cluster at the centre of the plot (Fig. 13), i.e. closest to the ‘average’ composition of historical Lamongan basalts. Significantly, Al2O3, Ni, Cr and Nb appear to exert little influence on the chemistry of erupted lavas, except possibly through minor variations in Plag content, suggesting that the abundance of these variables (i.e. high Al2O3, low Ni, Cr, Nb) was controlled primarily by the ‘source’ or in this case the magma supplied to the chamber.
A corollary of using PCA is that lavas of similar overall composition, and by inference of similar age, will cluster together in PC space. This is of use at Lamongan, where there is poor age control on several samples. The plot of PC3 against PC1 (Fig. 13) splits the samples into a ‘Y’-shaped array, with the 1821 samples grouping with sample LAM1207 on the lower ‘arm’. This latter sample was thus probably erupted at a similar time to the former, but from a different vent (Fig. 3). Similarly, samples LAM1261 (1847 flow) and LAM1563 (1821 flow?) also have similar characteristics (e.g. relatively high Nb contents) and hence may in fact be part of the same flow field (probably 1847; Fig. 3).
The cinder cone samples form a broadly L-shaped array in the plot of the first two PCs (Fig. 14), and the first three PCs account for 87% of the variance. This L-shaped pattern is often observed for suites of lavas from individual volcanoes, with picritic and differentiated flows occupying separate ‘arms’ and common compositions clustering at the apex (e.g. Albarède et al., 1997). In Fig. 14, one limb is composed of the KVT cones from the western LVF (west of the roughly north–south lineament defined by the three lake-filled maars immediately east of Klakah; Fig. 2) and two transitional high-K–shoshonitic samples (GTE288 and GAP1070). The KVM series and KVT samples from the central and eastern LVF form the other limb, with the younger cones plotting towards the left-hand side (Fig. 14). There is no appreciable clustering of samples at the ‘apex’, indicating that the LVF cones constitute two distinct magma series with little compositional overlap between them. However, several of the lavas that do plot in the region of the apex (GRG14071/2, GCK2960, GKN3060) are derived from cones situated close to the boundary between the compositionally distinct cones of the western LVF and the rest of the complex (Fig. 2).
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The western LVF samples are predominantly controlled by K2O, P2O5 and the incompatible elements, whereas the spread of compositions in the other lavas correlates with a CaO–MgO–Fe2O3–Sc–V–Co component (Cpx or possibly amphibole) and a SiO2–Na2O–Al2O3–Sr component (Plag), the former predominating in the KVM samples (Fig. 14). As observed in the historical lavas, Ni and Cr are not strongly correlated with the ferromagnesian component and plot at the centre of the diagram, suggesting minimal influence of mantle phases (Ni-rich Ol, Cr-rich pyroxene) on the major compositional trends and reflecting the overall low Ni and Cr contents of the suite. These elements exert more influence in the PC3 plot (Fig. 14), which probably reflects minor variations as a result of Ol accumulation in some magmas. Also in the PC3 diagram, the Al2O3 axis plots towards the pyroxene component, indicating the significance of Al-rich pyroxene in these lavas.
The spatial variation of magma composition in the LVF is also subtly manifested in the PC3 plot (Fig. 14), wherein three distinct geographical groups of cinder cones can be recognized. One group collects most of the cones on the northern and southern flanks of the LTT edifice (
; Fig. 14), which show a strong correlation with Fe2O3 and TiO2. Another assemblage (ß) contains the samples from a narrow region to the west of LTT and east of the three lake-filled maars (GMA5070, GCK2961/2960/1071, GKG136, GKL256, GRG14071/2, GKN3060; Fig. 2). This acts as a ‘transition zone’ between the previous group and the third group (
), which contains the KVT samples from the western LVF. This last group seems to correlate with Sr and the incompatible elements, and possibly Ni and Cr.
The PLF samples also define an L-shaped array in PC space (Fig. 15), with the PLF1 group clearly distinguished. These lavas correlate with the axes for K2O, P2O5 and the incompatible elements in the PC1–PC2 plot, whereas the PLF2/3 trend is controlled by a ferromagnesian component (Ol, Cpx) and Plag (SiO2, Al2O3, Na2O, Sr). Thus the PLF1 samples, which come from localities in the western LVF (Fig. 2), have similar overall characteristics to the KVT cones in the same region of the field.
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Contrary to the historical lavas and cinder cones, the PLFs show a more robust correlation between Ni, Cr and MgO. This indicates that the Ni and Cr budget is primarily controlled by the accumulation of phenocrysts rich in compatible elements (e.g. Ol) in the PLFs, and by the source in the historical lavas and cinder cones.
To summarize, many latent geochemical features of the LVF lavas are revealed by PCA. Spatial heterogeneities in magma composition are particularly apparent, with a clear distinction evident between lavas from the eastern and western parts of the LVF, and transitional samples with intermediate compositions. Contents of Rb, Ba, Y, Zr, K, P, ±Nb are interrelated in all samples, implying that these elements behaved incompatibly during magma evolution. Compatibility of Sr in Plag is indicated by the general correlation of Sr with the Plag component on plots of the first two PCs. Similarly, the general lack of correspondence between Ni, Cr and the ferromagnesian phases (except in the PLFs) attests to the low concentrations of these elements in the LVF suite (Fig. 11), which are probably a feature of the source magmas.
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FRACTIONAL CRYSTALLIZATION IN THE LVF LAVAS |
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Major elements
The results of PCA suggest that fractional crystallization of Ol, Cpx, Plag and Fe–Ti oxides may be responsible for much of the compositional variation within individual lava suites in the LVF. Mass balance calculations (e.g. Bryan et al., 1969
) have been employed to investigate this further, using mineral analyses from this work (Tables 4–6) with the exception of apatite (from Wood et al., 1995). Following Defant & Nielsen (1990), we have used 0·1 as an upper bound for the sum of the squares of the residuals (
r2), as values greater than this point to the involvement of other processes, such as assimilation or magma mixing, in addition to crystal fractionation during magma evolution. Selected results of mass balance calculations are given in Table 8. Predicted mineral assemblages are generally consistent with the observed modal mineralogy (Table 3), despite the evidence for crystal disequilibrium and accumulation (e.g. Fig. 6). It is thus unlikely that magma evolution in the LVF involved pure fractional crystallization, implying the removal of fractionating phases from the evolving melt, but that it entailed some mixing of phenocrysts and less evolved liquids. This is also implied by the generally scattered data in the Harker variation diagrams (Fig. 11).
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Mass balance modelling reveals potential genetic links between the KVM cones on the northern and southern flanks of LTT (east of sample ALA1161; Fig. 2), the entire series of historical lavas from Lamongan, and recent products of Semeru (e.g. sample SEM1607 from the February 1994 pyroclastic flow deposits; Table 2). These samples form a continuous sequence related by crystallization of variable proportions of Cpx + Plag + Mag + Ol (Table 8, numbers 1–7 and 16). On the basis of these calculations, samples GRJ288 or ALA1161 would be likely candidates for ‘parental’ magmas for recent lavas from Lamongan and Semeru.
Orthopyroxene (Opx) may enter the assemblage in the more evolved samples (e.g. LAM1561, SEM1607), being a common constituent of recent lavas at Semeru (e.g. the 1941–1942 lava flow; Table 2). There is little evidence for significant removal of Opx from the LVF suite, and its occurrence is restricted to the unstable cores and rims of Cpx crystals in the more evolved samples.
No acceptable r2 values were obtainable between the recent lavas and the KVT cones in the western LVF, indicating either that the two magma series are not genetically linked or that other processes were involved in the genesis of the KVT cones. Good matches were often achieved for all elements except the alkalis (Na2O and K2O), suggesting that mixing with an alkali-enriched (metasomatic?) component may link the two series, although this has not been tested. However, the least evolved sample from the western LVF (GOO13072; Fig. 2) is an acceptable parental composition for most of the KVT cones in that area and also for recently erupted lavas from Bromo (sample BRO1, erupted in March 1995; Table 2) to the west of the LVF (e.g. Table 8, numbers 8–10). A possible problem with this model is that the fractionating assemblage involves considerable amounts of Cpx (Table 8), which is only sparsely present in the lavas.
Several other clusters of cinder cones appear related to a common parental magma. Lavas from the array of collinear vents to the NW of LTT (GTU298–GGI2661; Figs 2 and 3) can all be derived from sample GPG2660, collected from a prehistoric lava flow (e.g. Table 8, numbers 11–13). Morphological evidence suggests that these vents and flows are of different ages and were probably emplaced along a roughly NW–SE-aligned fissure that has also produced maar-forming eruptions (Fig. 2; Carn, 2000). However, their related chemistries could indicate a contemporaneous, fissure-style eruption event, although the progressive exploitation over time of an eruptive fissure supplied by an evolving magma chamber would also be a feasible scenario.
The group of cones (KVM and KVT) to the west of the historical lava flow field on the flanks of Lamongan and east of the three lake-filled maars (Fig. 2) is also compositionally interrelated (e.g. Table 8, numbers 17 and 18). These samples, which were ‘transitional’ in the PCA diagram (Fig. 14), cannot, however, be related to any of the other cone clusters by a simple fractional crystallization process. As for the KVT cones in the western LVF, alkali contents are the main stumbling block in the mass balance calculations, and magma mixing may have been involved.
Calculations on the PLF samples suggest genetic links within and between the PLF2 and PLF3 series, involving relatively high proportions of Ol in the fractionating assemblage (e.g. Table 8, numbers 14 and 15). Fractionation between the PLF1 and PLF2/3 samples could not be successfully modelled, in accord with the orthogonal relationship between these two groups in PC space (Fig. 15).
Trace elements
Fractionation of trace elements was modelled using the Rayleigh fractionation equation, CL = COF(D – 1); where CL and CO are the concentrations of a given element in the daughter and parental liquids, respectively, F is the proportion of residual liquid and D is the bulk crystal–liquid distribution coefficient for that element (Shaw, 1970). Appropriate elemental partition coefficients (Kd values) were obtained from the literature (Table 9), with the crystallizing phases and proportions determined from the mass balance calculations (Table 8).
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This simple model generally reproduced the trace element contents of the modelled daughter liquids to an acceptable degree (Table 9). The only element that was consistently overestimated was V; this suggests that the Kd used for V in Cpx (0·8) may have been too low. Measured Kd values for V in amphibole are high (6 for basalts; e.g. Vukadinovic & Sutawidjaja, 1995), but there is minimal petrographic evidence for amphibole fractionation in these lavas.
Fractional crystallization is thus a potentially viable process to explain the major and trace element variations in separate series of LVF lavas. Plag and/or Cpx dominate most of the calculated assemblages, with Ol fractionation apparently significant in the evolution of the PLFs (Table 8). The crystallization sequence in the evolving magmas is difficult to pinpoint using petrographic evidence. On the basis of observed inclusion relationships, phenocryst sizes and degree of disequilibrium, a likely succession in the historical lavas and KVM samples would be Ol followed by Cpx and/or Mag, followed by Plag, late Opx and a late phase of Mag growth (or pyroxene breakdown) to produce the observed rims on Cpx phenocrysts. Both the KVT and PLF2 rocks show evidence for crystallization of Plag before Cpx.
With the exception of the KVT and PLF1 samples, the LVF lavas are generally highly phyric, with widespread evidence for phenocryst disequilibrium (e.g. Fig. 6). These observations raise the question that the compositions of some LVF lavas could be the product of essentially random mixing of phenocrysts and liquid, and hence bear little relation to a true liquid line of descent. However, the regular decrease in SiO2 (and increase in MgO) observed in the sequence of 19th-century lava flows erupted from Lamongan (Fig. 16) argues against such random sampling, at least in the historical lavas.
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SPATIAL GEOCHEMICAL VARIATIONS |
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Compositions of lavas erupted from cinder cones in the LVF appear to vary systematically with vent location, as documented in the preceding sections. Possible origins of this variation have been explored using incompatible element ratios, which are unaffected by fractional crystallization.
In Fig. 17, Zr/Y and Zr/K ratios are used to discriminate between the various geochemically distinct cinder cone clusters. Zr and Y are both relatively immobile in aqueous fluids (e.g. Tatsumi et al., 1986) and hence can provide information on the mantle source region of the lavas without contamination by fluids from the subducting slab. In the Zr/Y plot, the data bunch in a similar fashion to that observed in the PCA diagram (Fig. 14), although with slightly more overlap between groups. There is a clear separation between the KVT cones in the western LVF at high (>2·5) Zr/Y, and the eastern LVF cones at low (<2) values, with a ‘transitional’ group falling at intermediate ratios. The two shoshonitic samples also have high Zr/Y.
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; includes samples with prefixes GRG, GCK, GKG, GKL, GMA, GKN; i.e. group ß in Fig.
A possible explanation for the observed variation in Zr/Y ratios across the LVF is that the parental magmas for each cone group formed at different depths in the mantle wedge. Small degrees of melting at high pressures (>15 kbar) could result in elevated Zr/Y owing to retention of Y in garnet and Cpx (e.g. Tatsumi & Eggins, 1995), which would suggest a deeper source for the cones in the western LVF and the shoshonitic lavas (Fig. 17a). This would be consistent with the elevated K2O contents of these samples, as it is well known that K2O in erupted lavas often correlates positively with depth to the subducting slab in arc settings (e.g. Dickinson, 1975). What is not clear, however, is why magmas apparently generated at different depths should have erupted at similar distances from the arc front in the LVF.
Ratios of immobile to mobile incompatible trace elements (e.g. Zr/K) are more difficult to interpret, because of the effects of metasomatism of the mantle source on the mobile components. Zr/K ratios in the LVF cinder cones, however, show similar patterns to Zr/Y ratios, albeit with a few notable differences (Fig. 17b). As in the Zr/Y plot, the data cluster in spatial groups with the KVT cones in the western LVF displaying the highest ratios, but six of the ‘transitional’ samples and the shoshonitic rocks are displaced to lower Zr/K than the cones from the eastern LVF. This could be interpreted as the effect of addition of a K-rich or metasomatic component (e.g. slab-derived fluids; Vukadinovic & Nicholls, 1989) to the samples with low Zr/K, or that these lavas were produced by the melting of a metasomatic layer in the mantle beneath Java. If so, then Zr/K variations in the other LVF lavas may be due to varying degrees of metasomatism of their respective source regions, with the KVT cones in the western LVF associated with the least metasomatized mantle source.
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HOW PRIMITIVE ARE THE PARENTAL MAGMAS IN THE LVF? |
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According to the results presented above, there are at least two distinct LVF lava series: an older high-K suite (the KVT cones in the western LVF and the PLF1 series) and a calc-alkaline or medium-K suite (the KVM cones and historical lavas). A third series (GMA5070, GKG136, GCK2961, etc.; Fig. 2) possibly resulted from the mixing of these two end-member compositions and was erupted from vents situated spatially between them. The LVF high-K and medium-K series display genetic lineage with the recent lavas of Bromo and Semeru–Lamongan, respectively, and hence have the potential to contribute significantly to our understanding of magma genesis in the region.
Mass balance calculations (Table 8 and 9) suggest that samples GRJ288 and ALA1161 are close to parental compositions (i.e. the most primitive lavas recognized in the suite; Woodhead, 1988) for the medium-K series but, as discussed below, they are unlikely to represent true primitive or primary magmas. These samples are derived from two of the morphologically youngest prehistoric vents in the LVF (Carn, 2000), on the northern flanks of LTT (Fig. 2). Several features of these lavas indicate that they closely approximate true liquid compositions. Relative to many other LVF lavas, they are phenocryst poor (e.g. Table 3) and show few disequilibrium textures in the phenocrysts present (e.g. Table 7). In the case of sample ALA1161, the narrow range of Ol compositions plots within the high-pressure equilibrium field if it is assumed that Fe2O3/FeO = 0·2 in the host liquid (Fig. 6), indicating a probable absence of xenocrystic or accumulative Ol. However, the modelled parental magma for the high-K suite (GOO13071) probably has accumulated Ol (Fig. 6) and also exhibits some phenocryst disequilibrium textures (e.g. Table 4 and 6). In Table 10 the compositions of these parental LVF magmas are compared with analyses of primitive lavas from Pacific Rim arc volcanoes.
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Using the criteria of Tatsumi & Eggins (1995), who expected primary arc magmas that last equilibrated with peridotite in the mantle wedge to have high mg-numbers (
70) and high abundances of compatible trace elements (Ni > 200 ppm, Cr > 400 ppm), none of the LVF lavas, including the parental ones, can be realistically classed as primary. Their mg-numbers (<62), Ni (<22 ppm) and Cr (<25 ppm) contents are too low (Table 10), although the mg-number is highly dependent on the value of Fe3+/Fe2+ used. Two of the PLF2 samples have Ni and Cr contents of 60 and 150 ppm, respectively, and higher mg-numbers (71; Table 2), but this is probably a result of Ol accumulation in these lavas.
The discrepancy between primitive and parental magma compositions is typically attributed to the prior removal of olivine and/or pyroxene from the latter (e.g. Smith et al., 1997). This seems feasible for the more evolved parental magma (GOO13071; Table 10), but in the case of samples ALA1161 and GRJ288, their low SiO2 contents would seem to preclude any prior fractionation. It may, therefore, be necessary to invoke crystal accumulation or unusual source and/or melting conditions to explain the peculiar chemistry of these samples.
In addition to low SiO2, Ni and Cr contents, these lavas also have high TiO2, Al2O3, Fe2O3, CaO and low MgO relative to primitive magmas from circum-Pacific arcs (Table 10). Many lavas erupted in arc settings have high Al2O3, and they are of sufficient volumetric significance to have spawned the generic term high-Al2O3 basalts (HAB). Formal definitions of HAB composition vary, from aphyric basalts with Al2O3 >17 wt % (e.g. Tatsumi & Eggins, 1995), to lavas with <54 wt % SiO2 and >16·5 wt % Al2O3 (Crawford et al., 1987), and most of the LVF suite satisfy the latter criterion with the exception of several KVT and PLF1/2 samples. However, none of the LVF lavas can be described as aphyric, and indeed, the accumulation of plagioclase phenocrysts in evolved magmas is widely regarded to be the main process of HAB genesis, owing to a significant correlation between Al2O3 and modal plagioclase contents in arc volcanic rocks (e.g. Crawford et al., 1987). The modal data for the LVF suite are too limited to identify the presence or absence of any similar correlation (Fig. 18), but it appears that although elevated Al2O3 in many of the historical and KVM lavas is clearly related to abundant (>25%) modal Plag, the composition of sample ALA1161 (9% modal Plag) approaches that of a true HAB liquid. However, it could equally be construed as lying within the outer bounds of Crawford et al.’s (1987) correlation, so further investigation using REE data and Eu systematics (e.g. Woodhead, 1988), which can discriminate between fractionation and accumulation of plagioclase, would be required to debate this further.
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The low SiO2 contents of samples ALA1161 and GRJ288, and of many other LVF lavas, could also be construed as the result of phenocryst accumulation in more evolved liquids (e.g. Woodhead, 1988). However, if such processes can be discounted, which is possible in the case of sample ALA1161, then other reasons must be sought for the SiO2 deficiency, in addition to high CaO and Fe2O3 (Table 10). Various major elements, including Si, Fe and Ca, have shown potential as thermometers and barometers in basalts from ocean ridge and arc environments (e.g. Klein & Langmuir, 1987; Plank & Langmuir, 1993). In arc scenarios, high degrees of melting beneath thin arc crust (which allows a more extensive mantle column to undergo decompression melting, assuming melt initiation at a constant depth) are predicted to produce magmas with low Na, and high Ca and Fe (Plank & Langmuir, 1993). This is consistent with the chemistry of the parental LVF lavas (ALA1161 and GRJ288) and thus may be a viable model for magma genesis in the region. Using the expressions of Albarède [1992, equations (2) and (3)], we obtain a temperature of 1220 ± 40°C and a pressure of 15·5 ± 2·7 kbar for a primary mantle melt with SiO2 and MgO contents equivalent to ALA1161, although this is based on fluid-absent melting experiments. These conditions are similar to those obtained experimentally by Tatsumi et al. (1983) for the segregation of arc HAB from its source (15 kbar, 1340°C under anhydrous conditions; 17 kbar, 1325°C with 1·5 wt % H2O in the melt). Such temperatures are postulated to be excessively high for a stable mantle geotherm in a subduction zone, and are believed to demonstrate the existence of mantle diapirs beneath some arc volcanoes (e.g. Tatsumi et al., 1983).
Contents of V, and to a lesser extent Sc, in samples ALA1161 and GRJ288 are also high in comparison with primitive arc magmas (Table 10), although the low SiO2 contents of the LVF lavas make intercomparison difficult. V and Sc are particularly compatible in clinopyroxene, magnetite and amphibole, but the latter is considered to have the strongest affinity for these elements in basaltic liquids (e.g. Vukadinovic & Sutawidjaja, 1995). Although the effects of clinopyroxene accumulation cannot be ruled out, an alternative explanation for elevated V and Sc in these lavas is that the source region was abnormally rich in pyroxene or amphibole. Sisson & Bronto (1998) have recently inferred the involvement of pyroxenitic rocks in magma genesis at Galunggung volcano in West Java, in conjunction with a mantle diapir model similar to that propounded by Tatsumi et al. (1983). Low-degree melts of these pyroxene-rich lithologies entrained in upwelling mantle produce SiO2-undersaturated liquids with low H2O concentrations (Sisson & Bronto, 1998).
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DISCUSSION |
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Probably the most enigmatic aspect of the LVF is the occurrence of parental lavas, related both to the back-arc volcano, Bromo, and the arc-front volcano, Semeru, at similar distances from the trench. Furthermore, although Lamongan is in roughly the same position with respect to the trench as Bromo (Fig. 1), the historical and youngest prehistoric products of the volcano are medium-K basalts and basaltic andesites that can be genetically related to lavas currently erupting from Semeru. It is therefore likely that the two volcanoes are supplied by a common parental magma, of a composition similar to ALA1161 or GRJ288, and that these parental liquids are erupted as cinder cones in the LVF as a result of peculiar regional tectonics (see below).
The origin of the high-K cinder cones in the western LVF is unclear, but given their parental relationship to lavas from Bromo, it is possible that they represent the products of volcanism on the periphery of the Bromo–Tengger complex. These cones are amongst the oldest in the LVF and hence may not be temporally associated with activity in the rest of the field. However, geochronological work is needed to fully establish the age relationships between the various groups of vents identified in this study.
A working model for magma genesis beneath the LVF can be proposed, based on the salient geochemical features of the erupted lavas. It is stressed that this is probably only one of several possible models, but it is one that follows recent research on magma genesis at Javanese volcanoes (Sisson & Bronto, 1998) and that is broadly consistent with the LVF data. In this model, mantle upwelling and pressure-release melting occurs beneath the LVF. If extensional tectonics predominates in this part of East Java (as suggested by recent seismic crises in the LVF), the crust may be locally relatively thin, and thus high melt fractions can be achieved through the melting of hot, upwelling mantle. This produces primary HAB magmas with low SiO2 (if melting begins at high pressures) and Na2O and high CaO, FeO and Al2O3, which fractionate Ol, Cpx and calcic Plag during ascent to produce low-Mg HAB. Upon reaching the base of the crust small batches of this magma exploit extensional fractures and penetrate to the surface to erupt as cinder cones (e.g. ALA1161) or possibly maars; the rest may pond and fractionate further, or intrude into the crust and supply a small crustal magma chamber beneath the main volcanic edifice. The location of magma chamber formation, and hence of the LTT edifice in the LVF, may be controlled by the intersection of regional faults (Carn, 2000). Evolution in the magma chamber involves further fractionation of Ol, Cpx, Plag, Mag and minor Opx, along with the accumulation and resorption of phenocrysts. As fractionation proceeds, the residual liquid becomes increasingly volatile rich, until saturation is reached and exsolution of bubbles begins; this may trigger a phase of rapid growth of cellular Plag on pre-existing phenocrysts. Continuing exsolution increases the chamber pressure until eruption occurs, with crustal faults and fractures that intersect the chamber, possibly at deeper levels, acting as conduits for flank cinder cone and perhaps fissure-style eruptions (e.g. the KVM cones). The PLF2, PLF3 and historical lavas in the LVF are all interpreted as products of flank or summit eruptions from the LTT edifice supplied by a sub-volcanic chamber.
The KVT cones and PLF1 flows are assumed to represent an earlier phase of volcanism associated with the Tengger caldera, possibly before the initiation of regional extension and local thinning of the crust. This could explain the observed juxtaposition of ‘back-arc’ and ‘fore-arc’ magmas in the LVF. The chemical similarity between many of the maar deposits and the KVT–PLF1 samples (e.g. Fig. 10) suggests that the maar-forming eruptions often included fragments of the PLF1 flows as lithic material.
The compositionally distinct group of cones to the NE of Klakah (Fig. 2) could be the result of small-scale source heterogeneity, in conjunction with local fractures to convey the magmas to the surface whilst preserving their individuality. This is supported by the observation that these cones appear to vary considerably in age. Fracturing during the recent seismic crises in the LVF was concentrated in this region (Matahelumual, 1990). More detailed isotopic data from the LVF are needed to shed further light on this aspect of the complex. However, it is clear that spatial variations in source characteristics can be identified only where there is spatial variation in vent locations (i.e. in volcanic fields).
Although the regional tectonic framework of East Java is poorly known, it is clear that the LVF must occupy a zone of crustal extension. The existence of a monogenetic vent complex, the locally ‘pinched out’ topography of the island and the formation of the back-arc Madura Basin to the north (Fig. 1) all suggest an extensional regime. The LVF is also significant as it marks one of several locations on Java where the volcanic front is displaced away from the trench (Fig. 1). Gunung Slamet, a volcano in central Java, occupies a similar location and shares several structural characteristics with Lamongan (e.g. Vukadinovic & Sutawidjaja, 1995). Its central edifice also comprises two eruptive centres, split by a NW–SE-trending breach (similar to LTT; Figs 2 and 3), and it also has a scoria cone field on its flanks. These features may relate to the large-scale segmentation of the subduction system along this portion of the Sunda arc, with the LVF and Slamet situated at the boundaries of arc segments (Carn et al., in preparation).
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SUMMARY |
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Initial activity in the LVF may have involved the eruption of high-K back-arc basalts parental to the current products of the Bromo–Tengger system, in what is now the western part of the field. Subsequent or continuing extension perpendicular to the arc in East Java locally thinned the crust, and altered the composition of erupted lavas, producing low-Mg HAB. Exploitation of extensional fractures in the crust by small batches of rising magma led to the development of a cinder cone and maar field, and ultimately to the growth of a basaltic stratovolcano (LTT; estimated age
13–40 ka) at the intersection of two or more major lines of weakness in the crust. The magma chamber beneath LTT fed further cinder cone eruptions and crystal-rich lava flows on the immediate flanks of the volcano. However, the most primitive lavas exposed in the LVF are associated with some of the youngest prehistoric cinder cones, indicating relatively recent supply of new magma to the system. Extensional tectonics is still active in the LVF, as testified by the formation of ENE–WSW and east–west tensional fissures during the 20th-century seismic crises.
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ACKNOWLEDGEMENTS |
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We thank John Watson and Pete Webb at the Department of Earth Sciences, Open University, for assistance with XRF analyses. Hazel Chapman, Judith Bunbury, Maurice Haslop, Dave Newling and Stephen Reed (Cambridge) are gratefully acknowledged for guidance in Sr isotopic analysis, electron microprobe work and sample preparation. Fieldwork in Java was assisted by staff of the Volcanological Survey of Indonesia (VSI) from Bandung (Dedi Kusnadi), Yogyakarta (Dewi Subektiningsih) and Klakah (Sugeng Santosa, Bambang Wibono). We acknowledge helpful reviews by Ulrich Knittel and John Foden. S.A.C. was supported by a research studentship from the UK Natural Environment Research Council (GT4/94/137/G).
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FOOTNOTES |
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*Corresponding author. Present address: JCET, University of Maryland, Baltimore, M 21250, USA. E-mail: siman{at}jcet.umbc.edu
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REFERENCES |
|---|
Albarède, F. (1992). How deep do common basaltic magmas form and differentiate? Journal of Geophysical Research 97, 10997–11009.[Web of Science]
Albarède, F. (1995). Introduction to Geochemical Modelling. Cambridge: Cambridge University Press.
Albarède, F., Luais, B., Fitton, G., Semet, M., Kaminski, E., Upton, B. G. J., Bachèlery, P. & Cheminée, J.-L. (1997). The geochemical regimes of Piton de la Fournaise volcano (Réunion) during the last 530 000 years. Journal of Petrology 38, 171–201.
Apperson, K. D. (1991). Stress fields of the overriding plate at convergent margins and beneath active volcanic arcs. Science 254, 670–678.
Bacon, C. R., Bruggman, P. E., Christiansen, R. L., Clynne, M. A., Donnelly-Nolan, J. M. & Hildreth, W. (1997). Primitive magmas at five Cascade volcanic fields: melts from hot, heterogeneous sub-arc mantle. Canadian Mineralogist 35, 397–423.[Web of Science]
Borg, L. E., Clynne, M. A. & Bullen, T. D. (1997). The variable role of slab-derived fluids in the generation of a suite of primitive calc-alkaline lavas from the southernmost Cascades, California. Canadian Mineralogist 35, 425–452.[Web of Science]
Bronto, S., Situmorang, T. & Effendi, W. (1986). Geologic map of Lamongan volcano, Lumajang, East Java. Bandung: Volcanological Survey of Indonesia.
Bryan, W. B., Finger, L. W. & Chayes, F. (1969). Estimating proportions in petrographic mixing equations by least-squares approximation. Science 163, 926–927.
Campbell, I. H., Roeder, P. L. & Dixon, J. M. (1978). Plagioclase buoyancy in basaltic liquids as determined with a centrifuge furnace. Contributions to Mineralogy and Petrology 67, 369–377.
Camus, G., Gourgaud, A. & Vincent, P. M. (1987). Petrologic evolution of Krakatau (Indonesia): implications for a future activity. Journal of Volcanology and Geothermal Research 33, 299–316.[Web of Science]
Carn, S. A. (1999). Application of synthetic aperture radar (SAR) imagery to volcano mapping in the humid tropics: a case study in East Java, Indonesia. Bulletin of Volcanology 61, 92–105.[Web of Science]
Carn, S. A. (2000). The Lamongan volcanic field, East Java, Indonesia: physical volcanology, historic activity and hazards. Journal of Volcanology and Geothermal Research 95, 81–108.
Charlton, T. R. (1991). Postcollision extension in arc–continent collision zones, eastern Indonesia. Geology 19, 28–31.
Chesner, C. A. (1998). Petrogenesis of the Toba Tuffs, Sumatra, Indonesia. Journal of Petrology 39, 397–438.
Chotin, P., Giret, A., Rampnoux, J.-P., Sumarso, & Suminta (1980). L’île de Java, un enregistreur des mouvements tectoniques à l’aplomb d’une zone de subduction. Compte Rendu Sommaire des Séances de la Société Géologique de France 5, 175–177.
Crawford, A. J., Falloon, T. J. & Eggins, S. (1987). The origin of island arc high-alumina basalts. Contributions to Mineralogy and Petrology 97, 417–430.
Defant, M. J. & Nielsen, R. L. (1990). Interpretation of open system petrogenetic processes: phase equilibria constraints on magma evolution. Geochimica et Cosmochimica Acta 54, 87–102.
Dickinson, W. R. (1975). Potassium–depth (K–h) relations in continental margin and intraoceanic magmatic arcs. Geology 3, 53–56.
Dunn, T. & Sen, C. (1994). Mineral/matrix partition-coefficients for orthopyroxene, plagioclase, and olivine in basaltic to andesitic systems—a combined analytical and experimental study. Geochimica et Cosmochimica Acta 58, 717–733.[Web of Science]
Edwards, C. M. H., Morris, J. D. & Thirlwall, M. F. (1993). Separating mantle from slab signatures in arc lavas using B/Be and radiogenic isotope systematics. Nature 362, 530–533.
Edwards, C. M. H., Menzies, M. A., Thirlwall, M. F., Morris, J. D., Leeman, W. P. & Harmon, R. S. (1994). The transition to potassic alkaline volcanism in island arcs: the Ringgit–Beser complex, East Java, Indonesia. Journal of Petrology 35, 1557–1595.
Foden, J. D. (1983). The petrology of the calc-alkaline lavas of Rindjani volcano, east Sunda Arc: a model for island arc petrogenesis. Journal of Petrology 24, 98–130.[Web of Science]
Foden, J. D. & Varne, R. (1980). The petrology and tectonic setting of Quaternary–Recent volcanic centres of Lombok and Sumbawa, Sunda Arc. Chemical Geology 30, 201–226.
Gerbe, M.-C., Gourgaud, A., Sigmarsson, O., Harmon, R. S., Joron, J.-L. & Provost, A. (1992). Mineralogical and geochemical evolution of the 1982–1983 Galunggung eruption (Indonesia). Bulletin of Volcanology 54, 284–298.
Gust, D. A., Arculus, R. J. & Kersting, A. B. (1997). Aspects of magma sources and processes in the Honshu arc. Canadian Mineralogist 35, 347–365.[Web of Science]
Hamilton, W. (1979). Tectonics of the Indonesian Region. US Geological Survey Professional Paper 1078.
Hamilton, W. B. (1988). Plate tectonics and island arcs. Geological Society of America Bulletin 100, 1503–1527.
Hutchison, C. S. (1982). Indonesia. In: Thorpe, R. S. (ed.) Andesites: Orogenic Andesites and Related Rocks. New York: John Wiley, pp. 207–224.
Irvine, T. N. & Baragar, W. R. A. (1971). A guide to the chemical classification of the common volcanic rocks. Canadian Journal of Earth Sciences 8, 523–548.[Web of Science]
Jeffries, T. E., Perkins, W. T. & Pearce, N. J. G. (1995). Measurements of trace-elements in basalts and their phenocrysts by laser probe microanalysis inductively-coupled plasma-mass spectrometry (LPMA-ICP-MS). Chemical Geology 121, 131–144.
Jurewicz, A. J. G. & Watson, E. B. (1988). Cations in olivine, part 1: calcium partitioning and calcium–magnesium distribution between olivines and coexisting melts, with petrologic applications. Contributions to Mineralogy and Petrology 99, 176–185.
Katili, J. A. (1975). Volcanism and plate tectonics in the Indonesian island arcs. Tectonophysics 26, 165–188.
Klein, E. M. & Langmuir, C. H. (1987). Global correlations of ocean ridge basalt chemistry with axial depth and crustal thickness. Journal of Geophysical Research 92, 8089–8115.
Knittel, U. & Oles, D. (1995). Basaltic volcanism associated with extensional tectonics in the Taiwan–Luzon island arc: evidence for non-depleted sources and subduction zone enrichment. In: Smellie, J. L. (ed.) Volcanism Associated with Extension at Consuming Plate Margins. Geological Society, London, Special Publications 81, 77–93.
Knittel, U., Hegner, E., Bau, M. & Satir, M. (1997). Enrichment processes in the sub-arc mantle: a Sr–Nd–Pb isotopic and REE study of primitive arc basalts from the Philippines. Canadian Mineralogist 35, 327–346.
Kuno, H. (1968). Differentiation of basalt magmas. In: Hess, H. H. & Poldervaart, A. (eds) Basalts: the Poldervaart Treatise on Rocks of Basaltic Composition, Vol. 2. New York: Interscience, pp. 623–688.
Le Bas, M. J., Le Maitre, R. W., Streckeisen, A. & Zanettin, B. (1986). A chemical classification of volcanic rocks based on the Total Alkali–Silica diagram. Journal of Petrology 27, 745–750.
Le Maitre, R. W. (1968). Chemical variation within and between volcanic rock series—a statistical approach. Journal of Petrology 9, 220–252.
Leterrier, J., Yuwono, Y. S., Soeria-Atmadja, R. & Maury, R. C. (1990). Potassic volcanism in Central Java and South Sulawesi, Indonesia. Journal of Southeast Asian Earth Sciences 4, 171–187.
Lindsley, D. H. (1983). Pyroxene thermometry. American Mineralogist 68, 477–493.[Abstract]
Luhr, J. F. (1997). Extensional tectonics and the diverse primitive volcanic rocks in the western Mexican volcanic belt. Canadian Mineralogist 35, 473–500.[Web of Science]
Luhr, J. F. & Carmichael, I. S. E. (1981). The Colima volcanic complex, Mexico: Part II. Late Quaternary cinder cones. Contributions to Mineralogy and Petrology 76, 127–147.
Mandeville, C. W., Carey, S. & Sigurdsson, H. (1996). Magma mixing, fractional crystallization and volatile degassing during the 1883 eruption of Krakatau volcano, Indonesia. Journal of Volcanology and Geothermal Research 74, 243–274.
Matahelumual, J. (1990). Gunung Lamongan. Berita Berkala Vulkanologi, Edisi Khusus No. 125 (in Indonesian). Bandung: Volcanological Survey of Indonesia.
Morrice, M. G. & Gill, J. B. (1986). Spatial patterns in the mineralogy of island arc magma series: Sangihe arc, Indonesia. Journal of Volcanology and Geothermal Research 29, 311–353.
Mulyana, R. (1989). Pétrologie et géochimie du volcan Lamongan, Java–Indonesie. DEA thesis, Université de Bretagne Occidentale, Brest.
Nicholls, I. A. & Whitford, D. J. (1976). Primary magmas associated with Quaternary volcanism in the western Sunda Arc, Indonesia. In: Johnson, R. W. (ed.) Volcanism in Australasia. Amsterdam: Elsevier, pp. 77–90.
Nicholls, I. A. & Whitford, D. J. (1983). Potassium-rich volcanic rocks of the Muriah complex, Java, Indonesia: products of multiple magma sources? Journal of Volcanology and Geothermal Research 18, 337–359.
Nye, C. J. & Reid, M. R. (1986). Geochemistry of primary and least fractionated lavas from Okmok volcano, Central Aleutians: implications for arc magmagenesis. Journal of Geophysical Research 91, 10271–10287.
Panjasawatwong, Y., Danyushevsky, L. V., Crawford, A. J. & Harris, K. L. (1995). An experimental study of the effects of melt composition on plagioclase–melt equilibria at 5 and 10 kbar: implications for the origin of magmatic high-An plagioclase. Contributions to Mineralogy and Petrology 118, 420–432.
Pearce, J. A. & Peate, D. W. (1995). Tectonic implications of the composition of volcanic arc magmas. Annual Review of Earth and Planetary Sciences 23, 251–285.[Web of Science]
Peccerillo, A. & Taylor, S. R. (1976). Geochemistry of Eocene calc-alkaline volcanic rocks from the Kastamonu area, northern Turkey. Contributions to Mineralogy and Petrology 58, 63–81.[Web of Science]
Plank, T. & Langmuir, C. H. (1993). Tracing trace elements from sediment input to volcanic output at subduction zones. Nature 362, 739–743.
Puspito, N. T. & Shimazaki, K. (1995). Mantle structure and seismotectonics of the Sunda and Banda arcs. Tectonophysics 251, 215–228.
Rangin, C., Jolivet, L., Pubellier, M., Azéma, J., Briais, A., Chotin, P., Fontaine, H., Huchon, P., Maury, R., Müller, C., Rampnoux, J. P., Stephan, J. F. & Tournon, J. (1990). A simple model for the tectonic evolution of southeast Asia and Indonesia region for the past 43 m.y. Bulletin de la Société Géologique de France 8, 889–905.
Rickwood, P. C. (1989). Boundary lines within petrological diagrams which use oxides of major and minor elements. Lithos 22, 247–263.[Web of Science]
Roeder, P. L. & Emslie, R. F. (1970). Olivine–liquid equilibrium. Contributions to Mineralogy and Petrology 29, 275–289.[Web of Science]
Said, H. (1992). Gunung Lamongan. Berita Berkala Vulkanologi, Edisi Khusus No. 175 (in Indonesian). Bandung: Volcanological Survey of Indonesia.
Self, S., Rampino, M. R., Newton, M. S. & Wolff, J. A. (1984). Volcanological study of the great Tambora eruption of 1815. Geology 12, 659–663.
Self, S. & King, A. J. (1996). Petrology and sulfur and chlorine emissions of the 1963 eruption of Gunung Agung, Bali, Indonesia. Bulletin of Volcanology 58, 263–285.
Seyler, M. & Bonatti, E. (1994). Na, AlIV and AlVI in clinopyroxenes of subcontinental and suboceanic ridge peridotites: a clue to different melting processes in the mantle? Earth and Planetary Science Letters 122, 281–289.
Shaw, D. M. (1970). Trace element fractionation during anatexis. Geochimica et Cosmochimica Acta 34, 237–243.[Web of Science]
Silver, E. A., Reed, D. & McCaffrey, R. (1983). Back arc thrusting in the eastern Sunda arc, Indonesia: a consequence of arc–continent collision. Journal of Geophysical Research 88, 7429–7448.
Simkin, T. & Siebert, L. (1994). Volcanoes of the World, 2nd edn. Tucson, AZ: Geoscience Press.
Sisson, T. W. & Bronto, S. (1998). Evidence for pressure-release melting beneath magmatic arcs from basalt at Galunggung, Indonesia. Nature 391, 883–886.
Skogby, H. (1994). OH incorporation in synthetic clinopyroxene. American Mineralogist 79, 240–249.[Abstract]
Smith, I. E. M., Worthington, T. J., Price, R. C. & Gamble, J. A. (1997). Primitive magmas in arc-type volcanic associations: examples from the Southwest Pacific. Canadian Mineralogist 35, 257–273.[Web of Science]
Soeria-Atmadja, R., Maury, R. C., Bellon, H., Pringgoprawiro, H., Polve, M. & Priad, B. (1994). Tertiary magmatic belts in Java. Journal of Southeast Asian Earth Sciences 9, 13–27.
Stolz, A. J., Varne, R., Wheller, G. E., Foden, J. D. & Abbott, M. J. (1988). The geochemistry and petrogenesis of K-rich alkaline volcanics from the Batu Tara volcano, eastern Sunda Arc. Contributions to Mineralogy and Petrology 98, 374–389.
Stolz, A. J., Varne, R., Davies, G. R., Wheller, G. E. & Foden, J. D. (1990). Magma source components in an arc–continent collision zone: the Flores–Lembata sector, Sunda Arc, Indonesia. Contributions to Mineralogy and Petrology 105, 585–601.
Sun, S.-S. & McDonough, W. F. (1989). Chemical and isotopic systematics of oceanic basalts: implications for mantle compositions and processes. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications 42, 313–345.
Tatsumi, Y. & Eggins, S. (1995). Subduction Zone Magmatism. Cambridge, MA: Blackwell Science.
Tatsumi, Y., Sakuyama, M., Fukuyama, H. & Kushiro, I. (1983). Generation of arc basalt magmas and thermal structure of the mantle wedge in subduction zones. Journal of Geophysical Research 88, 5815–5825.[Web of Science]
Tatsumi, Y., Hamilton, D. L. & Nesbitt, R. W. (1986). Chemical characteristics of fluid phase released from a subducted lithosphere and the origin of arc magmas: evidence from high pressure experiments and natural rocks. Journal of Volcanology and Geothermal Research 29, 293–309.[Web of Science]
Tsuchiyama, A. (1985). Dissolution kinetics of plagioclase in the melt of the system diopside–albite–anorthite, and origin of dusty plagioclase in andesites. Contributions to Mineralogy and Petrology 89, 1–16.
Ulmer, P. (1989). The dependence of the Fe2+–Mg cation-partitioning between olivine and basaltic liquid on pressure, temperature and composition. Contributions to Mineralogy and Petrology 101, 261–273.
Van Bergen, M. J., Erfan, R. D., Sriwana, T., Suharyono, K., Poorter, R. P. E., Varekamp, J. C., Vroon, P. Z. & Wirakusumah, A. D. (1989). Spatial geochemical variations of arc volcanism around the Banda Sea. Netherlands Journal of Sea Research 24, 313–322.[Web of Science]
Van Gerven, M. & Pichler, H. (1995). Some aspects of the volcanology and geochemistry of the Tengger Caldera, Java, Indonesia: eruption of a K-rich tholeiitic series. Journal of Southeast Asian Earth Sciences 11, 125–133.
Varekamp, J. C., Van Bergen, M. J., Vroon, P. Z., Poorter, R. P. E., Wirakusumah, A. D., Erfan, R. D., Suharyono, K. & Sriwana, T. (1989). Volcanism and tectonics in the eastern Sunda arc, Indonesia. Netherlands Journal of Sea Research 24, 303–312.[Web of Science]
Varne, R. (1985). Ancient subcontinental mantle: a source for K-rich orogenic volcanics. Geology 13, 405–408.
Villemant, B., Jaffrezic, H., Joron, J. L. & Treuil, M. (1981). Distribution coefficients of major and trace-elements—fractional crystallization in the alkali basalt series of Chaîne-des-Puys (Massif Central, France). Geochimica et Cosmochimica Acta 45, 1997–2016.[Web of Science]
Vukadinovic, D. (1993). Are Sr enrichments in arc basalts due to plagioclase accumulation? Geology 21, 611–614.
Vukadinovic, D. (1995). High-field-strength elements in Javanese arc basalts and chemical layering in the mantle wedge. Mineralogy and Petrology 55, 293–308.[Web of Science]
Vukadinovic, D. & Nicholls, I. A. (1989). The petrogenesis of island arc basalts from Gunung Slamet volcano, Indonesia: trace element and 87Sr/86Sr constraints. Geochimica et Cosmochimica Acta 53, 2349–2363.
Vukadinovic, D. & Sutawidjaja, I. (1995). Geology, mineralogy and magma evolution of Gunung Slamet volcano, Java, Indonesia. Journal of Southeast Asian Earth Sciences 11, 135–164.[Web of Science]
Wahyudin, D. (1991). Volcanology and petrology of Mt. Semeru volcanic complex, East Java–Indonesia. M.Sc thesis, Victoria University, Wellington.
Wheller, G. E. & Varne, R. (1986). Genesis of dacitic magmatism at Batur volcano, Bali, Indonesia: implications for the origins of stratovolcano calderas. Journal of Volcanology and Geothermal Research 28, 363–378.
Wheller, G. E., Varne, R., Foden, J. D. & Abbott, M. J. (1987). Geochemistry of Quaternary volcanism in the Sunda–Banda Arc, Indonesia, and three-component genesis of island arc basaltic magmas. Journal of Volcanology and Geothermal Research 32, 137–160.
Whitford, D. J. (1975). Strontium isotopic studies of the volcanic rocks of the Sunda Arc, Indonesia, and their petrogenetic implications. Geochimica et Cosmochimica Acta 39, 1287–1302.
Whitford, D. J. & Nicholls, I. A. (1976). Potassium variations in lavas across the Sunda arc in Java and Bali. In: Johnson, R. W. (ed.) Volcanism in Australasia. Amsterdam: Elsevier, pp. 63–75.
Whitford, D. J., Nicholls, I. A. & Taylor, S. R. (1979). Spatial variations in the geochemistry of Quaternary lavas across the Sunda Arc in Java and Bali. Contributions to Mineralogy and Petrology 70, 341–356.
Whitford, D. J., White, W. M. & Jezek, P. A. (1981). Neodymium isotopic composition of Quaternary island arc lavas from Indonesia. Geochimica et Cosmochimica Acta 45, 989–995.
Widiyantoro, S. & van der Hilst, R. (1996). Structure and evolution of lithospheric slab beneath the Sunda arc, Indonesia. Science 271, 1566–1570.[Abstract]
Widiyantoro, S. & van der Hilst, R. (1997). Mantle structure beneath Indonesia inferred from high-resolution tomographic imaging. Geophysical Journal International 130, 167–182.[Web of Science]
Wood, C. P., Nairn, I. A., Mckee, C. O. & Talai, B. (1995). Petrology of the Rabaul Caldera area, Papua New Guinea. Journal of Volcanology and Geothermal Research 69, 285–302.[Web of Science]
Woodhead, J. D. (1988). The origin of geochemical variations in Mariana lavas: a general model for petrogenesis in intra-oceanic island arcs? Journal of Petrology 29, 805–830.
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