The Cretaceous Okhotsk–Chukotka Volcanic Belt (NE Russia): Geology, geochronology, magma output rates, and implications on the genesis of silicic LIPs

Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32

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Review

The Cretaceous Okhotsk–Chukotka Volcanic Belt (NE Russia): Geology, geochronology, magma output rates, and implications on the genesis of silicic LIPs P.L. Tikhomirov a, b,⁎, E.A. Kalinina a, T. Moriguti a, A. Makishima a, K. Kobayashi a, I.Yu. Cherepanova c, E. Nakamura a a b c

Institute for Study of the Earth's Interior, Okayama University at Misasa, Yamada 827, Misasa, Tottori 682-0193, Japan Geologic Faculty of Moscow State University, Leninskie Gory, Moscow 119991, Russia Chaun Mine Geologic Enterprise, Obrucheva 27, Pevek 686610, Russia

a r t i c l e

i n f o

Article history: Received 1 June 2011 Accepted 23 December 2011 Available online 3 January 2012 Keywords: Northeastern Eurasia Okhotsk–Chukotka belt Volcanism Geochronology Magma output rate Silicic LIP

a b s t r a c t The Cretaceous Okhotsk–Chukotka volcanic belt (OCVB) is a prominent subduction-related magmatic province, having the remarkably high proportion of silicic rocks (ca. 53% of the present-day crop area, and presumably over 70% of the total volcanic volume). Its estimated total extrusive volume ranges between 5.5 × 105 km3 (the most conservative estimate) and over 106 km3. This article presents a brief outline of the geology of OCVB, yet poorly described in international scientific literature, and results of a geochronological study on the northern part of the volcanic belt. On the base of new and published U–Pb and 40Ar/39Ar age determinations, a new chronological model is proposed. Our study indicates that the activity of the volcanic belt was highly discontinuous and comprised at least five main episodes at 106–98 Ma, 94–91 Ma, 89–87 Ma, 85.5–84 Ma, and 82–79 Ma. The new data allow a semi-quantitative estimate of the volcanic output rate for the observed part of the OCVB (area and volume approximately 105 km2 and 2.5 × 105 km3, respectively). The average extrusion rate for the entire lifetime of the volcanic belt ranges between 1.6 and 3.6 × 10− 5 km3yr− 1 km− 1, depending on the assumed average thickness of the volcanic pile; the optimal value is 2.6 × 10− 5 km3yr− 1 km− 1. Despite imprecise, such estimates infer the time-averaged volcanic productivity of the OCVB is similar to that of silicic LIPs and most active recent subduction-related volcanic areas of the Earth. However, the most extensive volcanic flare-ups at 89–87 and 85.5-84 Ma had higher rates of over 9.0 × 10− 5 km3yr− 1 km− 1. The main volumetric, temporal and compositional parameters of the OCVB are similar to those of silicic LIPs. This gives ground for discussion about the geodynamic setting of the latters, because the widely accepted definition of a LIP implies a strictly intraplate environment. Considering the genesis of the OCVB and other large provinces of silicic volcanism, we propose that residual thermal energy preserved in the continental crust after a previous major magmatic event may have been one of major reasons for high proportion of felsic rocks in a volcanic pile. In this scenario, underplating of mantle-derived basalts causes fast and extensive melting of still hot continental crust and generation of voluminous silicic magmas. © 2012 Elsevier B.V. All rights reserved.

Contents 1. 2. 3. 4. 5. 6. 7. 8.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . Geodynamic background . . . . . . . . . . . . . . . . . . Geologic description of the OCVB . . . . . . . . . . . . . . Age of the OCVB: a review of published data . . . . . . . . . The northern part of the OCVB: the area of a detailed study . . Petrographic description of samples studied . . . . . . . . . Analytical technique . . . . . . . . . . . . . . . . . . . . Results and discussion . . . . . . . . . . . . . . . . . . . 8.1. Volcanic episodes and spatial migration of OCVB activity 8.2. Volume of volcanic rocks and extrusion rates . . . . .

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⁎ Corresponding author at: Geologic Faculty of Moscow State University, Leninskie Gory, Moscow 119991, Russia. Tel./fax: + 7 495 9393865. E-mail addresses: [email protected] (P.L. Tikhomirov), [email protected] (E.A. Kalinina), [email protected] (T. Moriguti), [email protected] (A. Makishima), [email protected] (K. Kobayashi), [email protected] (E. Nakamura). 0377-0273/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2011.12.011

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8.3. OCVB and silicic LIPs: a semantic discussion . . . . 8.4. Origin of the OCVB and implications on the genesis 9. Summary . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . .

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1. Introduction The 3200 km-long Okhotsk–Chukotka volcanic belt (OCVB) of NE Eurasia extends from the western coast of the Sea of Okhotsk to the east of the Chukchi Peninsula (Fig. 1). With its present-day surface outcrop area over 450,000 km 2, the OCVB appears to be the largest volcanic province spatially related to active continental margins. It constitutes a significant part of the Mesozoic magmatic arc system of the Circum-Pacific (e.g., Nokleberg et al., 2001) and comprises over 1 million km3 of volcanic rocks, not including its eroded parts (Belyi, 1977; Kotlyar et al., 1981). Among other subduction-related volcanic belts, the OCVB has an unusually high proportion of silicic rocks, up to 80–85 vol.% in some segments. The OCVB has been studied by geologists since 1930-s (e.g. Obrutchev, 1934). An interest to the belt is additionally supported by significant economic mineral resources (gold, silver, tin, and mercury). Sparcity of post-magmatic deformation and relatively weak rock alteration, as

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well as variably advanced erosional incision (200–300 m to 2–3 km), make the OCVB particularly attractive for petrologic and paleovolcanologic studies. However, there is a serious lack of precise analytical data, like quantitative analyses for trace elements and isotopic analyses. Several monographs (e.g., Belyi, 1977, 1994; Kotlyar et al., 1981; Filatova, 1988; Struzhkov and Konstantinov, 2005) and numerous papers dedicated to the OCVB have been published in Russian; mainly, these publications are focused on stratigraphy and mineral resources of the volcanic belt. Papers published in international journals are somewhat sparse, and usually comprise data on relatively small localities within this large volcanic province (e.g., Kelley et al., 1999; Polin and Moll-Stalcup, 1999; Ispolatov et al., 2004; Hourigan and Akinin, 2004; Tikhomirov et al., 2006, 2008; Mishin et al., 2008; Stone et al., 2009; Sakhno et al., 2010). At the moment, the most representative set of geochronological and geochemical data on the OCVB obtained by precise methods have been published by Akinin and Miller (2011). However, the latest English-language paper which contains more or less detailed

Fig. 1. The position of the Okhotsk–Chukotka volcanic belt in the tectonic frame of NE Asia (compiled and modified after Tilman and Bogdanov, 1992; Nokleberg et al., 2001). 1 — Siberian craton and minor cratonal blocks; 2 — complexes of the Paleozoic to Early Mesozoic passive margin of the Siberian continent; 3 — the ‘Kolyma–Omolon superterrane’, the tectonic collage of terranes of various nature (including remnants of Paleozoic to Mesozoic island arcs, backarc, forearc and intra-arc basins, accretionary wedges, passive margins, and minor cratonal blocks), amalgamated before the Early Cretaceous; 4 — remnants of Late Jurassic to Early Cretaceous volcanic arcs; 5 — South Anyui suture zone; 6 — partially eroded late orogenic Tytylveyem volcanic belt (Tikhomirov et al., 2009a); 7 — syn-collisional basins filled by Late Jurassic to Early Cretaceous clastic sediments; 8 — Magadan and South Taigonos batholiths of granitic rocks, 9 — the Okhotsk–Chukotka volcanic belt (OCVB): a — the boundary between its frontal and rear zones, b — boundaries between its segments, after Belyi (1977) (characters depict the names of segments: WOS — West Okhotsk, OS — Okhotsk, PS — Penzhina, ANS — Anadyr, CCS — Central Chukotka, and ECS — East Chukotka); 10 — the proposed position of the trench during the OCVB activity; 11 — the Koryak–Kamchatka tectonic province (various terranes, related mainly with the Late Cretaceous to Cenozoic island arcs, accreted during the Cenozoic era); 12 — Cenozoic continental basins filled with clastic sediments; 13 — outlines of the area depicted on Fig. 2.

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general geologic description of the OCVB has been published in 1960-s (Belyi et al., 1966). Taking into account the changes in overall viewpoints on global tectonic processes and the accumulation of a large body of new data, the revised version of the geologic overview of the OCVB appears appropriate. The large volcanic volume and high proportion of felsic rocks make the OCVB similar to a specific class of large igneous provinces (LIPs) called silicic LIPs, or SLIPs (Bryan et al., 2002). The similarity between these provinces and the OCVB looks intriguing because the most recent definition of LIPs proposed by Bryan and Ernst (2008) implies the formation of 'true' LIPs (i.e. those which fit the specified volumetric and temporal criteria) in intraplate settings. The comparison of the key characteristics of the OCVB with those of SLIPs may become an important argument in the discussion on the tectonic prerequisites of a SLIP formation, and help understanding of the mechanisms controlling the large-volume silicic volcanism in general. The main obstacle which hampers such a comparison is lack of reliable geochronological data for the OCVB. The problems outlined above define the structure of this paper which includes 3 main parts: (1) a brief review of the general geological information on the OCVB; (2) the analysis of the present-day knowledge of the OCVB geochronology and presentation of the new geochronological model built on both published ( 40Ar/ 39Ar, U–Pb) and new (U–Pb) isotopic dates from the northern OCVB, and (3) comparison of the OCVB with SLIPs and discussion on some genetic aspects of large-volume silicic volcanism.

2. Geodynamic background Formation of the OCVB postdated amalgamation of several compositionally and geodynamically diverse terranes including cratonal blocks, Late Paleozoic to Early Cretaceous island arcs, accretionary and passive margin complexes, and minor ophiolites; and their accretion to the Siberian continent (Parfenov, 1991; Nokleberg et al., 2001). The present-day position of major terranes of NE Asia is shown on Fig. 1. At the final stage of the formation of this tectonic collage, Arctic Alaska–Chukotka microcontinent has collided with the active margin of Siberia, and the Anyui–Angayucham oceanic basin closed to form the South Anyui suture zone (Sokolov et al., 2002). Several syn-collisional basins (Figs. 1 and 2) are filled by weakly deformed Late Jurassic through Early Cretaceous clastic, locally coal-bearing sediments, up to 3 km thick. These basins commonly record gradual transition from marine to continental sedimentation (Belyi, 1994; Filatova, 1988). The maximum age of last major compressional event may be inferred from deformations of Early Cretaceous (up to 127 Ma) strata in the South Anyui suture zone (Sokolov et al., 2002). The Late Jurassic–Early Cretaceous collisions terminated the earlier arc volcanism at the margins of the Oimyakon and Anyui–Angayucham oceans (Parfenov, 1991; Nokleberg et al., 2001). The direct precursor of the OCVB is the Late Jurassic to Early Cretaceous Uda–Murgal arc formed at the active margin of ancient Pacific ocean (Sokolov et al., 2009). The fragments of this arc are present in the basement of the OCVB along its almost entire length, usually under its frontal zone

Fig. 2. Simplified geologic map of the Western and Central Chukotka area (after Varlamova et al., 2004, modified). 1 — uplifts of the Paleozoic basement; 2 — Permian(?) to Triassic clastic sedimentary complexes of the passive margin of Chukotka microcontinent; 3 — sedimentary and volcanic complexes of the Oloy zone (remnants of backarc and forearc basins of Jurassic to Early Cretaceous age); 4 — remnants of Late Jurassic to Early Cretaceous island arcs; 5 — areas of late orogenic volcanism of Aptian and Albian age (characters in squares: Tt — Tytylveyem, N — Nutesyn, and L — Lyadindya volcanic depressions); 6 — syn-collisional basins filled with clastic sediments of Late Jurassic to Early Cretaceous, sometimes coal-bearing (circled characters: A — Ainakhkurgen, R — Rauchua depressions); 7–13 — the Okhotsk–Chukotka volcanic belt. Tick marks depict the area used for the calculations of volumetric characteristics and volcanic output rates (see Section 8.2). Different types of tick marks correspond to products of distinct volcanic episodes: 7–8 — 106–98 Ma (7 — mafic to intermediate (the ‘lower andesites’), 8 — silicic); 9 — 94–91 Ma, silicic; 10 — 89–87 Ma, mainly silicic; 11–85.5–84 Ma, mainly silicic; 12–79–74(?) Ma, dominantly mafic (the ‘upper basalts’); 13 — the part of the OCVB excluded from volumetric calculations; 14 — depressions filled with Cenozoic clastic sediments; 15 — granitic and gabbro-granitic plutons (Paleozoic to Late Cretaceous); 16 — mafic and ultramafic rocks of ophiolitic complexes; 17 — inferred former outlines of the partially eroded late orogenic Tytylveyem volcanic belt of the Aptian age; 18 — the conventional boundary between the rear and the frontal zones of the OCVB (after Belyi, 1977); 19 — the boundaries between OCVB segments (ANS — Anadyr, CCS — Central Chukotka, ECS — East Chukotka segments); 20 — the position of the transect of the OCVB (see Section 6). Star symbols depict the sampling points for zircon U–Pb dating (open stars correspond to the published data of Moll-Stalcup et al., 1995; Tikhomirov et al., 2006; Miller et al., 2009; Sakhno et al., 2010; filled symbols depict the samples used for this study). Open triangles depict the sampling points for the 40Ar/39Ar dating (Kelley et al., 1999; Ispolatov et al., 2004; Tikhomirov et al., 2006; Sakhno et al., 2010). The appropriate isotopic ages of volcanic rocks are presented with 2σ error values. Shaded rectangles depict areas with stratigraphic columns presented in Fig. 3.

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(Fig. 1). Conventionally, the Uda–Murgal volcanic belt is thought to be a continental arc (Parfenov, 1991; Sokolov et al., 2001), however some interpretations imply its northern segments evolved as an island arc (Filatova, 1988; Morozov, 2001) so there was a transition from continental to island arc from south to north (Lawver et al., 2002). The youngest formations of the Uda–Murgal arc are thought to have the Barremian age (127 to 121 Ma; Varlamova et al., 2004), and the accretion of the Pekulney arc presumably took place at the same time (Morozov, 2001). After the extinction of Jurassic to Early Cretaceous arcs but before the onset of the OCVB activity, some post-orogenic volcanism took place in local zones discordant to the paleo-Pacific margin (Kotlyar and Rusakova, 2004; Tikhomirov et al., 2009a). Within the area of the accreted Chukotka microcontinent, several chains of post-collisional granite batholiths were emplaced during the Early Cretaceous (Varlamova et al., 2004; Miller et al., 2009; Tikhomirov et al., 2011). The OCVB activity is thought to be related to the northwestward subduction of Kula and/or Isanagi oceanic plates (Parfenov, 1991; Nokleberg et al., 2001; Ueda and Miyashita, 2005) under the newly formed composite Verkhoyansk–Chukotka orogen. Volcanic sequences with total thickness up to 6 km (Belyi, 1977) unconformably overlie the heterogeneous Mesozoic tectonic collage. The position of the paleo-trench is still ambiguous. Gently folded clastic sediments (up to 3 km thick) of a forearc basin crop near the eastern margin of the volcanic belt, but the only proved occurrence of the coeval accretionary complex has been reported for the Omgon Range, West Kamchatka (Soloviev et al., 2006), about 350 km east from the OCVB axis (Fig. 1). One may surmise the accretion process had a minor significance during the OCVB activity, or the greater part of the accretionary wedge has been removed by later tectonic processes, or it is just buried under Late Cenozoic sediments of the Penzhina–Anadyr basin (Fig. 1). In any case, the estimated distance between the trench and the OCVB volcanic centres would range from 100 to 500 km. This implies likely shallow-dipping subduction zone, comparable to this of the recent Andean active margin (e.g., Tatsumi and Eggins, 1995). At the end of Cretaceous, a major re-arrangement of plate kinematics (e.g.,

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Cox et al., 1989) caused eastward shift of the subduction zone and the cessation of subduction-related volcanism within the OCVB. 3. Geologic description of the OCVB The basement of the OCVB is quite heterogenous (Fig. 1). It includes several cratonic blocks and various complexes of both passive and active continental margins, with rare fragments of ophiolitic sections (e.g., Nokleberg et al., 2001; Sokolov et al., 2002). The rear (continent-side) margin of the OCVB looks strongly discordant to major fold axes, faults, and boundaries of tectonic zones of the basement (Fig. 1). The frontal (ocean-side) margin of the volcanic belt unconformably overlies remnants of older volcanic arcs of ancestral Pacific, of nearly the same trend as the OCVB. Sediments of syn-collisional basins, coeval with the Uda–Murgal arc, may show folding and thrusting (Baranov, 1996; Miller and Verzhbitsky, 2009), but sometimes they are virtually undeformed, and their contact with the OCVB base may look conformable. Local post-orogenic volcanic formations (for example, those of the 121–112 Ma Tytylveyem belt — Tikhomirov et al., 2009a) occupy similar structural position, being only slightly deformed. It is commonly thought that the OCVB includes two major zones (e.g., Belyi, 1994; Filatova, 1988; Sidorov et al., 2009), the 'inner' (frontal, ocean-side) and the ‘outer’ (rear, continent-side) ones (Figs. 1 and 3), which have diffrerent predominant rock composition and total thicknesses of a volcanic pile. The ocean-side zone consists mostly of andesites and basaltic andesites, and comprises thicker volcanic section of 4–6 km, up to 7.5 km (Sidorov et al., 2009). The rear zone is thinner (2–5 km), and comprises a greater proportion of silicic rocks, varying from 10 to 85 percent along strike of the belt (Kotlyar et al., 1981; Varlamova et al., 2004). However, the significant part of the frontal zone defined by Belyi (1977, 1994) corresponds to the outcrops of the Uda–Murgal arc, which has been proven to represent a separate, earlier stage of the history of Pacific active margin. (e.g., Parfenov, 1991; Filatova, 1988). After ‘subtraction’ of Uda–Murgal arc complexes from the OCVB, the average thickness of the OCVB and the proportion of felsic rocks appear to approach these of the rear zone. The cross-arc

Fig. 3. Generalized stratigraphic columns of the OCVB (modified after Filatova, 1988; the segments of the OCVB after Belyi, 1977). 1–2 — sediments of syn-collisional basins (1 — marine, 2 — continental); 3–7 — main volcanic series of the OCVB, dominated by rhyolites (3), rhyolites and dacites (4), andesites (5), basalts (6) and alkaline basalts (7); 8 — main unconformities and disconformities.

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compositional zoning of the belt defined by decreasing proportion of felsic rocks towards the ocean still exists but is less pronounced. In some OCVB parts surface observations do not reveal any notable difference between volcanic complexes of rear and frontal zones (for example, in the northern part of the OCVB shown on Fig. 2). The data on the total thickness of the OCVB are quite ambiguous. Seismic data are available for a single transect north from the Magadan city (Simonov et al., 2003). These data suggest the thickness of a volcanic pile of up to 2 km, provided that the seismic profile intersects the Okhotsk segment of the OCVB, which is presumably thinner than other parts of the volcanic belt (see below). Conventionally, the stratigraphic columns for the OCVB (Figs. 3 and 4) have been built using estimated thicknesses of local units exposed at surface. Such estimates are too imprecise because of strong spatial variability of the volcanic pile. Both lateral dimensions of any unit concealed by younger strata, and variations of its thickness are quite difficult to be assessed. The calculations based on gravity field data do not provide a reliable result because the gravity field may be strongly affected by blind geologic bodies of an unknown size and shape (e.g., granitic plutons). Conventionally, the estimated average thickness of the OCVB ranges between 2 and 4 km (Filatova, 1988; Belyi, 1994; Sidorov et al., 2009). The additional discussion on these matters is presented in Section 8.2. In addition to cross-arc zoning, six transverse segments of the volcanic belt are being distinguished (Fig. 1). These segments differ from each other by various volume proportions of felsic and mafic/ intermediate rocks, as well as by some major element characteristics, like average alkali-silica ratio (Kotlyar et al., 1981). Boundaries between the segments are gradational rather than sharp. The Okhotsk segment may have a lesser total thickness, compared тo other parts of the volcanic belt. Here, the volcanic fields are somewhat discontinuous, with abundant exposures of basement within the volcanic belt. As evident from the generalized sections of main segments of the OCVB (Fig. 3; modified after Filatova, 1988), stratigraphy of the belt consists of three main components: (1) the ‘lower andesites’, (2) the group of formations dominated by silicic rocks, and (3) the ‘upper

basalts’. The latter are sometimes interpreted as unrelated to the OCVB, but linked to a later rifting event. (Filatova, 1988; Kotlyar and Rusakova, 2004). Minor beds of epiclastic rocks are scattered through the entire section of the OCVB. Limited spatial extent of these layers, absence of any marine fauna, and presence of fossil flora indicate that all major units of the OCVB have been formed in a continental environment. To the east and south, the volcanic rocks of the OCVB grade into thick (at least 2 km, and over 3 km in some locations) clastic sequences of the marine forearc basin. At some locations, the forearc sediments are overlain by Late Cretaceous continental volcanics up to 1.5 km thick, suggesting shift of volcanic activity towards the ocean during the late stages of OCVB evolution. Sequences of the OCVB do not reveal evidences of superimposed regional deformations. The exception is the frontal zone of the Penzhina segment of the volcanic belt, where several local thrusts have been mapped (Montin, 1992). Over most of the OCVB, dips of volcanic strata do not exceed 15°, being likely controlled by paleotopography. This is not the case in vicinity of calderas, which are quite ubiquitous. Near margins of some calderas, dips locally reach 80° and even 90° (Tikhomirov, 1996). In addition to calderas which normally do not exceed 20 km in diameter, there are gentle subsidence structures of a larger size, either linear, up to 150 km long (and concordant with the strike of the volcanic belt), or subequant of up to 100 km across. At the surface, their marginal parts are expressed by monoclines with dip angles of 5–15° (Fig. 5). The linear extension-related folds akin those described in the Sierra Madre Occidental province (Ferrari et al., 2002) have not been documented in the OCVB, and most researchers (e.g., Yarmolyuk, 1973; Belyi, 1977; Umitbaev, 1986; Filatova, 1988) relate the observed monoclines with a roof subsidence of large, likely mid-crustal magma chambers. The ‘lower andesitic’ units comprise basalts, basaltic andesites and andesites, with minor dacites, rhyolites, and epiclastic interlayers. The proportion between lavas and pyroclastic rocks is variable, but lavas are typically less abundant (30–40 vol.%). During this first stage of the OCVB formation, stratovolcanoes were the most common type of a volcanic edifice (Filatova, 1988). In some OCVB segments (e.g., Central Chukotka segment), the ‘lower andesites’ are sparse, and the

Fig. 4. Simplified stratigraphic columns for the studied part of the Okhotsk–Chukotka volcanic belt (after Varlamova et al., 2004). Each column corresponds to an appropriate location depicted on Fig. 2. 1 — folded sedimentary basement (Triassic); 2 — the Late Jurassic to Early Cretaceous clastic sediments; 3–7 — volcanic rocks, dominated by rhyolites (3), dacites (4), andesites (5), basalts (6) and shoshonitic rocks (7); 8 — shallow-level silicic intrusions sampled for isotopic dating; 9 — the most pronounced unconformities. All volcanic units contain minor epiclastic rocks which portion (1 to 5% in vol.) tends to increase towards the top of the total section of the OCVB. Thicknesses of volcanic suites correspond to their maximal known values, hence the total thickness of the volcanic pile may be overestimated. Stars and triangles mark the stratigraphic position of samples used for isotopic age determinations (see the caption for Fig. 2). The appropriate isotopic ages of volcanic rocks are given. For each newly obtained U–Pb date, the isotopic age is presented as numerator (with 2σ error values), and denominator indicates the sample number.

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Fig. 5. Satellite image (LANDSAT) of a volcanic monocline in the western Central Chukotka segment of the OCVB, SW from the El'gygytgyn lake (the location inside the shaded area c on Fig. 2). White marks with numbers depict dip and strike of volcanic strata, yellow dashed lines correspond to contacts between compositionally distinct units.

available volcanic record begins with silicic rocks. Basalts and andesites are porphyritic, with phenocrysts of plagioclase, clinopyroxene, orthopyroxene, Ti-magnetite, amphibole, and apatite. In basalts, minor olivine and accessory Cr-spinel may be present. The middle part of the OCVB section is dominated by dacites and rhyolites, mainly as ignimbrites and tuffs. Sometimes, silicic volcanics are intercalated with mafic rocks similar with those present in the lower units of the OCVB. The mafic/silicic ratio widely varies both in space and time. The proportion between silicic ignimbrites, ash-fall tuffs, lavas and epiclastic rocks is also variable. In general, abundance of ignimbrites tends to decrease upwards. Thick (1–3 km) sequences of intensely welded ignimbrites are usually confined to calderas. Some ignimbritic formations are poorly stratified, and spread over an area of n ⁎ 10 3 km 2 (Varlamova et al., 2004). Many geologists consider these units to result from a series of catastrophic pelean type eruptions (e.g., Umitbaev, 1986). Outflow deposits are better stratified; they usually comprise a greater portion of ash-fall tuffs and epiclastic rocks, and few relatively thin (up to 100–200 m) ignimbritc units. In silicic volcanics of the OCVB, the content of phenocrysts varies between 0 and 40–45%. The most ubiquitous phenocryst phase is plagioclase of a wide compositional range (An65-10), though the most calcic grains are likely xenocrysts. Amphibole, K–Na feldspar, biotite, quartz, and both pyroxenes may be also abundant. The accessory phases are Ti–magnetite, ilmenite, apatite, zircon, allanite, sometimes garnet, monazite, xenotime, and thorite. The glassy matrix is usually devitrified, unaltered obsidian like rocks (vitrophyres) are relatively sparse. These vitrophyres are dark-coloured, black and brown, contrasting with light-coloured devitrified felsic volcanics. The amount of fresh volcanic glass gradually increases towards the top of the OCVB section. The ‘upper basaltic’ formations are represented by basalts, trachybasalts, basaltic andesites, and basaltic trachyandesites. In presentday topography, they are commonly expressed by plateaus of up to 10,000 km 2. Rhyolitic ash-fall tuffs and ignimbrites locally intercalate with mafic rocks, and constitute to 10% of total volume of these units. Petrographically, the ‘upper basalts’ commonly differ from older mafic rocks of the OCVB. Some of them have a greater percentage of olivine phenocrysts, and some may contain titaniferous augite and/or minor analcime. On the contrary, felsic rocks intercalated with 'upper basalts' are similar to other silicic rocks of the OCVB. Published data on geochemistry of volcanic rocks of the OCVB are rather sparse, especially trace element data (Polin and Moll-Stalcup, 1999; Kalinina et al., 2008; Akinin and Miller, 2011). The ‘lower andesites’ display a clear arc affinity, having a well pronounced negative Ta–Nb anomaly, Pb spike, and high LILE/HFSE ratio. The ‘upper basalts’

19

may retain similar features, but frequently they are relatively enriched, and yield the characteristics transitional to those of intraplate basalts (Filatova, 1988; Kalinina et al., 2008). The felsic rocks are quite various in trace element and isotopic composition, suggestive for a heterogeneity of magma sources (Tikhomirov et al., 2008). The detailed work on the geochemistry of the OCVB is currently in progress; in this paper, we focus on temporal aspects of the OCVB evolution. Some features inherent to the OCVB (e.g., superposition on a relatively young continental crust, main temporal changes of rock composition, arc-to-intraplate geochemical affinity) are usual of many other provinces of silicic volcanism, including ‘recognized’ SLIPs (Bryan and Ernst, 2008) and smaller provinces having an eruptive volume of over 104 km3, for example, Late Paleozoic volcanic areas of Central and Western Europe (Wilson et al., 2004), Mid-Paleozoic Lachlan fold belt (Cas, 1983), Mid-Late Paleozoic belts of Kazakhstan (Bakhteev, 1987), Mesozoic Great Xing'an belt (Zhang et al., 2008), Tertiary Great Basin area (Lipman, 1992), recent Taupo zone (Houghton et al., 1995; Charlier et al., 2004), and others. This similarity offers a possibility to analyse factors which control the eruption of large-volume silicic magmas throughout the Earth. A brief discussion of these matters is presented in Section 8.4, though the thorough consideration of the problem is worth a separate publication. 4. Age of the OCVB: a review of published data The history of the OCVB is still being debated. The 'traditional' model based on paleobotanic correlations and whole rock K–Ar and Rb–Sr dates (e.g. Belyi, 1994; Kotlyar and Rusakova, 2004; Zhulanova et al., 2007) implies beginning of active eruptions in Albian (108 ± 3 Ma), and the cessation of volcanism in Campanian, at 75 ± 3 Ma. If the 'upper basalts' are assumed unrelated to the OCVB (e.g., Filatova, 1988), estimates of the upper time limit of suprasubductional volcanism range from early Turonian (92 ± 2 Ma; Kotlyar and Rusakova, 2004) to early Campanian (80 ± 3 Ma; Filatova, 1988). There are controversies regarding the age of some complexes of paleoflora (e.g., Belyi, 1994; Herman, 1999; Kotlyar and Rusakova, 2004, and references therein). Main factors limiting reliability of paleobotanic methods as stratigraphic correlation tool within the OCVB and other similar volcanic provinces, are: (1) relatively rare findings of diagnostic plant fossils; (2) the spatial variability of volcanic strata, hampering their lithologic correlations; and (3) commonly equivocal age correlations between the continental flora and marine fossils (e.g., Herman, 1999). The more detailed discussion on these matters is presented in Tikhomirov et al. (2006). The numerous K–Ar dates obtained for magmatic rocks of the OCVB yield a wide age interval, 50 to 110 Ma (Kotlyar and Rusakova, 2004), and frequently conflict with documented field relationships. Lack of confidence in these data does not allow using them as solid constraints on the age of the OCVB. However, some researchers consider the oldest K–Ar ages as reflecting the real age of the initial volcanic activity (Zhulanova et al., 2007). They conclude that both K–Ar data and bulk rock Rb–Sr isochrons (near 20) support the traditional chronological model that is based primarily on paleobotanic data. Relatively precise and reliable radiometric dating methods, such as 40Ar/ 39Ar dating of mineral separates (sanidine, amphibole, biotite, plagioclase) and SHRIMP U–Pb zircon dating, that became widely used since the mid 1990-s (Moll-Stalcup et al., 1995; Kelley et al., 1999; Ispolatov et al., 2004; Hourigan and Akinin, 2004; Tikhomirov et al., 2006; Mishin et al., 2008; Sakhno et al., 2010), frequently return dates that disagree with traditionally accepted ages of the OCVB units. Despite the overall lifetime of the OCVB outlined using 40Ar/ 39Ar and U–Pb determinations (106 ± 1.7 to 74 ± 1.2 Ma; Akinin and Miller, 2011) is almost the same with that obtained by paleobotanic methods, the timing of the most active eruptions has been re-assessed. The greater part of OCVB volcanic pile seems to be accumulated between 89 and 81 Ma (with two main statistical peaks at ca. 87 and 82 Ma),

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P.L. Tikhomirov et al. / Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32

whereas some previous models imply a lull period in OCVB activity during this time interval (e.g., Kotlyar and Rusakova, 2004). The 'upper basalts' yield ages of 78.8 ± 1.2 to 74 ± 1.2 Ma, which are not much younger than those for the uppermost silicic units and coeval granitic plutons (79.2 ± 1.9 Ma; Tikhomirov et al., 2006). The only sample from the ‘upper basaltic’ unit of the East Chukotka segment of the OCVB (Fig. 2) yields a much younger U–Pb age of 67 ± 1 Ma (Sakhno et al., 2010). This could result both from a spatial diachronism of the OCVB activity, or from a superposed volcanic event, related to the Cenozoic active margin of the Pacific. Hence, published 40Ar/ 39Ar and zircon U–Pb dates suggest that the timing of the OCVB as a whole generally agrees with the traditional model, but the chronostratigraphy of some its parts requires a revision. Besides, the old K–Ar and paleobotanic data do not allow to distinguish main episodes of the OCVB activity. Reconstruction of spatial migrations of volcanism also requires a more comprehensive geochronological study, with relatively dense sampling of an area of hundreds km across. 5. The northern part of the OCVB: the area of a detailed study At the moment, implementation of a research project covering entire 3200 km-long OCVB requires considerable funds and efforts. As the first step, we present results of our study of the northern part of the belt, which is better exposed (due to location in tundra climatic zone) and well explored during 1960-s to 1980-s. Samples for this study were collected in the Central Chukotka segment of the OCVB (CCS), in the northern part of the Anadyr segment (ANS), and in the western part of the East Chukotka segment (Fig. 2). Here, the volcanic belt abruptly changes its strike, and has a maximum width of up to 250 km. The total outcrop area of the part of OCVB presented on Fig. 2 is about 130,000 km 2. The area of nearly 90,000 km2 was used for volumetric calculations (see Section 8.2). Compared to central and southern segments of the OCVB, this area comprises a greater proportion of silicic rocks (70–75% vol. on average). Generalized stratigraphic columns of the northern OCVB are shown on Fig. 3, with names and thicknesses of units after Varlamova et al. (2004). Thicknesses of volcanic formations (suites) are maximum known thicknesses, hence the total thickness of the volcanic pile may be overestimated (for volumetric calculations, we assumed lesser values — see Section 8.2). Published 40Ar/39Ar and zircon U–Pb ages (Kelley et al., 1999; Ispolatov et al., 2004; Tikhomirov et al., 2006; Sakhno et al., 2010), and the newly obtained zircon U–Pb dates are also shown. On average, the volcanic strata of the CCS and ANS gently dip southeast. Only in the narrow southeastern marginal zone of a volcanic belt, the Late Cretaceous volcanics unconformably overlie the folded remnants of the Pekulney arc (Fig. 2), dipping west and northwest. 6. Petrographic description of samples studied Fig. 2 illustrates locations of 15 new U–Pb zircon samples, as well as sampling points for all published U–Pb and 40Ar/ 39Ar age data on northern OCVB. The stratigraphic position of samples is shown on Fig. 4, and their coordinates are listed in Table 1. All samples were collected from stratified volcanics, except sample 04–558 which represents a dike of an obsidian-like rhyolite, probably coeval with the youngest formations of the ANS rear zone. Together with the previously published 23 U–Pb and 40Ar/ 39Ar dates, geochronological dataset characterizes all major stratigraphic units of the area shown on Fig. 2. The main emphasis was put on dating of samples from the western part of CCS, where the rock samples were collected along a 270-km transect across the OCVB (Fig. 2), with the continuous documentation of petrographic composition and attitude of volcanic units. The studied part of the OCVB comprises volcanic rocks of a wide compositional range, including basalts, basaltic andesites, andesites,

shoshonites, latites, dacites, and rhyolites. As noted in Section 5, felsic rocks (rhyolites with minor dacites) are predominant. Pyroclastic rocks (tuffs and ignimbrites) constitute more than 90% of total volume of silicic rocks. For age determinations, we used samples of silicic porphyritic lavas or crystal-rich pyroclastic rocks with the negligible content of lithic clasts, to avoid xenocrysts. Among 15 samples analysed, there are 7 rhyolitic tuffs, 3 rhyolitic ignimbrites, 2 dacitic ignimbrites, 1 rhyolitic lava, 1 rhyolite from a dike, and 1 latite lava. The silicic tuffs and ignimbrites are composed of crystal and vitric material in various proportions. The most common phenocryst phases are feldspars: (1) plagioclase An55-11 in rhyolites and An60-22 in dacites; (2) sanidine Ab29–46Or51–69An1–4, frequently transformed into perthitic orthoclase or microcline by hydrothermal processes; and sometimes (3) anorthoclase Ab72–82Or4–20An5–19. Phenocrysts of quartz, biotite, amphibole, and hypersthene En46–74Fs23–52Wo1–3 are also ubiquitous. Relics of hedenbergite En43–48Fs13–15Wo39–42 occur in some dacites. Most common accessory phases are titaniferous magnetite, ilmenite, apatite, zircon, sphene, and allanite are abundant; whereas almandine-spessartine garnet, monazite, and thorite are less common. The matrix of felsic tuffs and ignimbrites is composed of variably welded glassy shards and lapilli-size vitric fragments. In most samples, silicic glass is replaced by felsitic or microgranular quartz–feldspar aggregate. Felsic lavas are less common than pyroclastic rocks. They have a massive, or flow-banded, or, rarely, amygdaloidal textures. Phenocrysts are the same as in pyroclastic rocks, but not fragmented. Total content of phenocrysts ranges from 0 to 12–15%. Groundmass consists of devitrified glass replaced by felsitic, granoblastic, poikiloblastic, or axiolitic quartz-feldspar aggregate, which sometimes includes microlites of plagioclase and minor K-feldspar, up to 5–7% in vol. Several silicic samples comprise fresh glass, either in lava matrix, or in ash particles, or in fiamme. Among felsic volcanics studied, there are rocks containing quartz and feldspar phenocrysts of comparable size (probably crystallized from magmas of a near-cotectic composition), and rocks in which quartz is minor or absent (likely derivatives of less silicic magmas). Volumetric proportions between these two rock groups vary widely. The latite 2582-1/460 of the Etchikun suite (Fig. 4a) is massive brownish grey porphyritic rock. Phenocrysts (15 to 25% vol.) comprise pyroxene up to 4 mm, plagioclase up to 3 mm, K–Na feldspar up to 5 mm, biotite (up to 3 mm, some in glomerocrysts with feldspars), rare prismatic zircon, abundant acicular apatite, titaniferous magnetite, and ilmenite. The former glassy matrix contains microlites of plagioclase and K-Na feldspar. Both groundmass and phenocrysts underwent low temperature alteration, being replaced by aggregates of albite, greenish chlorite, sericite, quartz, clay minerals, and leucoxene. Hydrothermal alteration assemblages are similar in all rock varieties, though modal proportions of secondary phases depend on rock composition. The modal abundance of hydrothermal minerals varies from sample to sample but generally increases with the stratigraphic age. Main epigenetic minerals are chlorites, albite, adularia, epidote, chalcedony, quartz, clay minerals, sericite, celadonite, stilpnomelane, carbonates, and zeolites. In a few thin sections, minor babingtonite, julgoldite, prehnite and pumpellyite were identified. Secondary minerals may replace the groundmass of volcanic rocks, or form pseudomorphs after phenocrysts, or fill cracks, amygdules and other voids. In all samples used for age determinations, the alteration is weak or absent.

7. Analytical technique Composition of mineral phases was analyzed using Scanning Electron Microscope Hitachi S-3100H with the Energy Dispersive X-ray spectrometer Horiba EMAX 7000 operating at accelerating voltage of 20 kV, beam current of 0.3 nA, and counting time of 100–110 s.

P.L. Tikhomirov et al. / Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32

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Table 1 SIMS U–Pb isotopic analyses of zircons from volcanic rocks of the northern OCVB. U (ppm)

Th (ppm)

Th/U

206

2582-1/460 2582-1/460-1 2582-1/460-2 2582-1/460-3 2582-1/460-11 2582-1/460-5a

17 16 10 16 13

10 9 7 10 8

0.56 0.61 0.64 0.63 0.62

1444 1404 1000 812 538

0.0494 ± 0.0011 0.0503 ± 0.0014 0.0491 ± 0.0011 0.0484 ± 0.0008 0.0525 ± 0.0013

0.1137 ± 0.0217 0.1144 ± 0.0247 0.1142 ± 0.0258 0.1158 ± 0.0165 0.1261 ± 0.0283

0.0165 ± .0011 0.0165 ± 0.0011 0.0169 ± 0.0012 0.0173 ± 0.0009 0.0169 ± 0.0006

104.4 ± 4.5 104.6 ± 5.4 108.0 ± 6.3 111.0 ± 4.0 106.1 ± 2.2

T05-31 T05-31-1a T05-31-4a T05-31-6 T05-31-9 T05-31-10 T05-31-20

231 344 179 406 195 657

124 104 78 64 58 424

0.54 0.30 0.44 0.16 0.30 0.65

6142 7479 4154 9079 4846 14971

0.0478 ± 0.0478 0.0486 ± 0.0486 0.0490 ± 0.0490 0.0477 ± 0.0477 0.0467 ± 0.0467 0.0473 ± 0.0473

0.1049 ± 0.0129 0.1119 ± 0.0112 0.1063 ± 0.0158 0.1069 ± 0.0098 0.0982 ± 0.0144 0.1018 ± 0.0087

0.0158 ± 0.0007 0.0167 ± 0.0005 0.0158 ± 0.0008 0.0161 ± 0.0006 0.0154 ± 0.0012 0.0155 ± 0.0008

100.7 ± 3.1 106.5 ± 2.3 100.7 ± 3.2 103.0 ± 2.2 99.4 ± 6.8 99.9 ± 4.1

04-418 04-418-1 04-418-2 04-418-3 04-418-6 04-418-7 04-418-10 04-418-11

528 656 306 803 379 1069 546

207 281 111 338 88 711 163

0.39 0.43 0.36 0.42 0.23 0.67 0.30

10685 12665 6297 19429 7948 24292 12311

0.0491 ± 0.0004 0.0488 ± 0.0004 0.0491 ± 0.0007 0.0488 ± 0.0004 0.0493 ± 0.0005 0.0481 ± 0.0004 0.0491 ± 0.0005

0.1053 ± 0.0075 0.1041 ± 0.0079 0.1059 ± 0.0121 0.1029 ± 0.0082 0.1042 ± 0.0115 0.1029 ± 0.0089 0.1033 ± 0.0090

0.0156 ± 0.0006 0.0154 ± 0.0005 0.0157 ± 0.0006 0.0153 ± 0.0007 0.0152 ± 0.0007 0.0155 ± 0.0004 0.0154 ± 0.0005

97.7 ± 2.4 97.8 ± 1.6 99.6 ± 2.5 96.2 ± 2.7 95.4 ± 2.5 98.9 ± 1.3 97.4 ± 1.8

T05-34 T05-34-5 T05-34-8 T05-34-8a T05-34-4

1178 709 775 476

350 306 357 202

0.30 0.43 0.46 0.42

12834 10355 13048 5870

0.0488 ± 0.0005 0.0483 ± 0.0006 0.0486 ± 0.0005 0.0479 ± 0.0007

0.0987 ± 0.0096 0.0992 ± 0.0169 0.0981 ± 0.0104 0.1002 ± 0.0132

0.0147 ± 0.0006 0.0150 ± 0.0013 0.0146 ± 0.0011 0.0151 ± 0.0009

93.7 ± 2.6 95.8 ± 7.3 92.4 ± 5.9 96.2 ± 5.1

2238a/02 2238a/02-1 2238a/02-2 2238a/02-4 2238a/02-5 2238a/02-6 2238a/02-7 2238a/02-8

159 367 737 413 624 314 188

63 172 164 262 254 166 105

0.40 0.47 0.22 0.64 0.41 0.53 0.56

3324 8808 15135 10525 13973 6278 4335

0.0485 ± 0.0008 0.0497 ± 0.0005 0.0485 ± 0.0004 0.0486 ± 0.0006 0.0489 ± 0.0004 0.0491 ± 0.0008 0.0485 ± 0.0007

0.0991 ± 0.0170 0.1034 ± 0.0118 0.0964 ± 0.0149 0.0928 ± 0.0117 0.0981 ± 0.0114 0.1002 ± 0.0150 0.1005 ± 0.0148

0.0148 ± 0.0012 0.0149 ± 0.0009 0.0143 ± 0.0018 0.0139 ± 0.0011 0.0145 ± 0.0012 0.0145 ± 0.0010 0.0151 ± 0.0011

94.0 ± 6.0 91.9 ± 3.7 89.0 ± 10.0 88.2 ± 5.3 89.9 ± 5.9 90.5 ± 4.3 96.0 ± 5.4

2320a/02 2320a/02-3 2320a/02-5 2320a/02-10 2320a/02-12 2320a/02-18 2320a/02-14

242 331 494 403 172 127

68 132 386 180 63 60

0.28 0.4 0.78 0.45 0.37 0.47

6073 8916 11560 9161 4004 3374

0.0490 ± 0.0006 0.0487 ± 0.0004 0.0494 ± 0.0004 0.0490 ± 0.0005 0.0488 ± 0.0006 0.0508 ± 0.0010

0.0987 ± 0.0127 0.1014 ± 0.0090 0.1021 ± 0.0097 0.0955 ± 0.0102 0.0996 ± 0.0166 0.0977 ± 0.0186

0.0144 ± 0.0008 0.0150 ± 0.0009 0.0149 ± 0.0007 0.0141 ± 0.0009 0.0146 ± 0.0010 0.0139 ± 0.0011

90.9 ± 3.9 94.0 ± 4.5 93.0 ± 2.8 88.6 ± 4.6 92.1 ± 4.4 87.0 ± 5.0

04-558 04-558-7 04-558-3 04-558-4 04-558-5 04-558-6

762 649 335 399 375

585 250 170 187 196

0.77 0.39 0.51 0.47 0.52

6331 5333 2803 3876 3512

0.0450 ± 0.0005 0.0462 ± 0.0006 0.0458 ± 0.0012 0.0449 ± 0.0010 0.0461 ± 0.0008

0.0838 ± 0.0096 0.0901 ± 0.0116 0.0839 ± 0.0181 0.0862 ± 0.0161 0.0863 ± 0.0153

0.0133 ± 0.0006 0.0140 ± 0.0005 0.0132 ± 0.0008 0.0138 ± 0.0006 0.0136 ± 0.0006

87.4 ± 2.4 90.3 ± 2.3 85.4 ± 4.6 89.4 ± 3.3 88.0 ± 2.9

T05-93-1 T05-93-1-1 T05-93-1-2 T05-93-1-5 T05-93-1-4 T05-93-1-14

168 254 413 118 606

92 154 404 69 535

0.55 0.60 0.98 0.58 0.88

2832 3075 7733 2141 11001

0.0478 ± 0.0010 0.0485 ± 0.0008 0.0475 ± 0.0006 0.0479 ± 0.0009 0.0479 ± 0.0005

0.0902 ± 0.0177 0.0946 ± 0.0115 0.0903 ± 0.0092 0.0914 ± 0.0142 0.0892 ± 0.0093

0.0136 ± 0.0011 0.0141 ± 0.0005 0.0138 ± 0.0007 0.0137 ± 0.0007 0.0134 ± 0.0008

86.6 ± 5.7 89.7 ± 2.4 88.5 ± 3.1 87.7 ± 3.5 84.9 ± 3.9

T06-22 T06-22-1 T06-22-2 T06-22-3 T06-22-4 T06-22-6 T06-22-7 T06-22-8

623 343 582 523 596 620 640

228 102 216 265 202 195 240

0.37 0.30 0.37 0.51 0.34 0.31 0.38

14487 8415 18696 12313 15429 14566 10975

0.0484 ± 0.0004 0.0486 ± 0.0005 0.0482 ± 0.0004 0.0484 ± 0.0005 0.0479 ± 0.0005 0.0477 ± 0.0004 0.0475 ± 0.0004

0.0933 ± 0.0086 0.0951 ± 0.0096 0.0926 ± 0.0105 0.0924 ± 0.0081 0.0914 ± 0.0087 0.0914 ± 0.009 0.0908 ± 0.0087

0.0139 ± 0.0006 0.0141 ± 0.0007 0.0139 ± 0.0009 0.0138 ± 0.0007 0.0138 ± 0.0010 0.0138 ± 0.0007 0.0139 ± 0.0008

88.2 ± 2.4 88.9 ± 3.1 88.1 ± 4.1 87.2 ± 2.9 88.1 ± 5.2 87.4 ± 2.6 89.0 ± 3.0

T05-37 T05-37-1 T05-37-2 T05-37-5

9 28 14

4 18 7

0.40 0.63 0.54

2161 5348 2990

0.0493 ± 0.0011 0.0498 ± 0.0006 0.0485 ± 0.0008

0.0923 ± 0.0186 0.0942 ± 0.0101 0.0884 ± 0.0145

0.0135 ± 0.0007 0.0136 ± 0.0004 0.0131 ± 0.0008

85.8 ± 3.6 85.6 ± 1.9 83.3 ± 3.9

Analysis N

Pb/204Pb

207

Pb/206Pb (± 2σ)

207

Pb/235U (± 2σ)

206

Pb/238U (± 2σ)

Age (Ma) (± 2σ)

(continued on next page)

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P.L. Tikhomirov et al. / Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32

Table 1 (continued) U (ppm)

Th (ppm)

Th/U

206

T05-37 T05-37-6 T05-37-8 T05-37-9

13 28 10

9 14 5

0.64 0.51 0.45

3660 4713 2923

0.0467 ± 0.0009 0.0479 ± 0.0007 0.0497 ± 0.0009

0.0884 ± 0.0161 0.0879 ± 0.0124 0.0906 ± 0.0148

0.0135 ± 0.0006 0.0133 ± 0.0007 0.0131 ± 0.0006

86.6 ± 2.6 85.3 ± 3.6 83.3 ± 3.3

T05-66 T05-66-2 T05-66-3 T05-66-8 T05-66-3a T05-66-9 T05-66-10 T05-66-11 T05-66-12 T05-66-13 T05-66-14

14 10 18 9 23 19 18 12 15 19

9 6 9 5 9 7 11 6 8 7

0.60 0.59 0.51 0.53 0.36 0.39 0.59 0.48 0.54 0.38

2986 1525 3924 1384 3919 3323 2965 1784 2940 2808

0.0479 ± 0.0009 0.0484 ± 0.0012 0.0492 ± 0.0009 0.0470 ± 0.0013 0.0493 ± 0.0007 0.0485 ± 0.0008 0.0484 ± 0.0007 0.0472 ± 0.0013 0.0505 ± 0.0010 0.0485 ± 0.0008

0.0900 ± 0.0144 0.0906 ± 0.0192 0.0949 ± 0.0149 0.0948 ± 0.0239 0.0910 ± 0.0115 0.0942 ± 0.0156 0.0949 ± 0.0135 0.0920 ± 0.0214 0.1067 ± 0.0181 0.0907 ± 0.0143

0.0134 ± 0.0007 0.0136 ± 0.0008 0.0142 ± 0.0006 0.0142 ± 0.0004 0.0133 ± 0.0003 0.0140 ± 0.0004 0.0140 ± 0.0004 0.0141 ± 0.0004 0.0152 ± 0.0005 0.0135 ± 0.0004

84.9 ± 3.3 87.2 ± 4.5 90.5 ± 2.9 90.9 ± 1.9 84.6 ± 1.1 89.7 ± 1.6 89.2 ± 1.2 90.5 ± 2.2 96.0 ± 1.8 85.8 ± 1.3

7155a/89 7155a/89-1 7155a/89-2 7155a/89-3 7155a/89-4

282 546 840 775

139 291 539 399

0.49 0.53 0.64 0.51

5603 9818 15145 14343

0.0499 ± 0.0007 0.0498 ± 0.0005 0.0481 ± 0.0004 0.0482 ± 0.0004

0.0942 ± 0.0116 0.0915 ± 0.0101 0.0881 ± 0.0060 0.0879 ± 0.0068

0.0136 ± 0.0004 0.0134 ± 0.0007 0.0133 ± 0.0003 0.0132 ± 0.0002

85.6 ± 1.4 83.9 ± 3.1 84.57 ± 0.86 84.32 ± 0.68

88-4824 88-4824-1 88-4824-2 88-4824-3 88-4824-4

16 11 19 14

8 5 9 10

0.51 0.46 0.47 0.75

650 2152 3612 2283

0.0490 ± 0.0009 0.0481 ± 0.0011 0.0482 ± 0.0006 0.0481 ± 0.0012

0.0921 ± 0.0145 0.0862 ± 0.0168 0.0918 ± 0.0100 0.0851 ± 0.0181

0.0134 ± 0.0006 0.0130 ± 0.0008 0.0137 ± 0.0007 0.0127 ± 0.0007

85.0 ± 3.0 83.1 ± 1.1 87.4 ± 1.1 81.1 ± 3.2

T05-86-1 T05-86-1-2 T05-86-1-3 T05-86-1-4 T05-86-1-5

177 143 331 250

187 136 455 320

1.06 0.95 1.38 1.28

199 239 404 920

0.0905 ± 0.0142 0.0926 ± 0.0175 0.0977 ± 0.0207 0.0965 ± 0.0268

0.0905 ± 0.0142 0.0926 ± 0.0175 0.0977 ± 0.0207 0.0965 ± 0.0268

0.0134 ± 0.0007 0.0134 ± 0.0011 0.0134 ± 0.0009 0.0132 ± 0.0009

85.3 ± 3.5 83.9 ± 5.9 83.0 ± 4.0 83.9 ± 5.3

T05-88 T05-88-9 T05-88-8 T05-88-8a T05-88-10 T05-88-12

589 226 181 198 424

390 181 170 214 379

0.66 0.8 0.94 1.08 0.89

6514 2858 2129 2090 3778

0.0865 ± 0.0127 0.0893 ± 0.0160 0.0878 ± 0.0164 0.0855 ± 0.0177 0.0850 ± 0.0132

0.0865 ± 0.0127 0.0893 ± 0.0160 0.0878 ± 0.0164 0.0855 ± 0.0177 0.0850 ± 0.0132

0.0130 ± 0.0008 0.0132 ± 0.0011 0.0133 ± 0.0008 0.0128 ± 0.0012 0.0127 ± 0.0012

83.0 ± 4.0 83.2 ± 5.9 85.0 ± 4.5 81.3 ± 6.8 80.0 ± 6.1

Analysis N

Pb/204Pb

207

Zircons for U–Pb age determinations were extracted by conventional separation techniques including crushing, magnetic and heavy liquid separation. Some analyses were undertaken using polished thin sections after their microscopic and SEM-EDX study. SEM S-3100H was used for the cathodoluminescent (CL) imaging. Examined grains are prismatic euhedral crystals 50 to 250 μm long. Acicular apatite inclusions are common (Fig. 6a–c), so transmitted light images were necessary to select an appropriate area for the analysis. CL images display an oscillatory zoning in all studied zircons from volcanic rocks, and no evidences of inherited cores (Fig. 6d–f). We used a high resolution secondary ion mass spectrometer (HRSIMS) Cameca IMS-1270 at the Pheasant Memorial Laboratory for Geochemistry and Cosmochemistry of the University of Okayama at Misasa, Japan. Analytical procedures are similar to those described by Usui et al. (2002). Zircons were sputtered with a focused O− primary beam of 5 or 15 nA intensity with 13.0 kV accelerating potential, resulting in a ~ 10 μm beam diameter. During the analysis, dynamic multicollection of secondary ions was done employing the HR-SIMS equipped with four movable and one fixed secondary electron multipliers. This reduces the data acquisition time and cancels the time drift of the secondary ion beam intensity caused by the instability of the primary ion beam and the charging effect on the sample. The signals for 204Pb +, 206Pb +, 207Pb + and 208Pb + were simultaneously detected by the individual multipliers appropriately configured on the focal plane. The magnet power was switched from Pb + to UO + and UO + signals were detected by the immovable centre multiplier. Instrumental mass discrimination and mass fractionation for U/Pb atomic ratios and Pb isotopic ratios were corrected by applying

Pb/206Pb (± 2σ)

207

Pb/235U (± 2σ)

206

Pb/238U (± 2σ)

Age (Ma) (± 2σ)

calibration curves directly obtained using the Geostandard Sri Lanka zircon with an age of 561 ± 2 Ma (1σ error) reported by Usui et al. (2002). Data reduction was performed using the Isoplot software (Ludwig, 2000). Crosschecking of the sputtered craters was done using a polarizing transmitted and reflected light microscope in order to confirm whether the probed pits incorporated inclusions and/or cracks. All age errors are reported at 2σ. 8. Results and discussion Results of zircon U–Pb dating are summarized in Table 1 and on Figs. 2, 3, 8, and 9. All analyzed samples yield concordant ages, well consistent with observed stratigraphic relationships and with the published 40Ar/ 39Ar and U–Pb age data (Kelley et al., 1999; Ispolatov et al., 2004; Tikhomirov et al., 2006; Sakhno et al., 2010), within the analytical error (usually, the 2σ value is 1 to 2 Ma). Among available 40Ar/ 39Ar ages, we consider only those obtained on mineral separates (sanidine, amphibole, biotite, and plagioclase). Bulk rock determinations appear to provide less reliable results, perhaps because of the poor ability of the groundmass of volcanic rocks to retain radiogenic argon. Indeed, the whole rock 40Ar/39Ar dates sometimes contradict the stratigraphic relationships (e.g., Stone et al., 2009), though most of them are reasonably consistent with zircon U–Pb and 40Ar/ 39Ar mineral separate dates. For most samples, the ages of individual zircon grains do not show any considerable scatter beyond the analytical error, thus suggesting the inherited Pb content is negligible, and no significant Pb loss took place. The only sample T05-66 includes zircons which may belong

P.L. Tikhomirov et al. / Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32

23

Fig. 6. The transmitted light (a-c) and cathodoluminescent (d-f) images of 3 typical zircons from volcanic rocks of the OCVB. Note the abundance of inclusions (apatite, minor feldspars and melt inclusions) and the absence of inherited cores.

to different age populations: (1) 89.78 ± 0.77 (5 grains, MSWD = 0.79), (2) 85.19 ±0.80 Ma (4 grains, MSWD = 0.73), and (3) one single grain with the U–Pb age of 96.8 ± 1.8 Ma (Fig. 7o and p). The group of zircons with the age of about 90 Ma has been interpreted as antecrysts (e.g., Charlier et al., 2004), and the youngest group as the products of crystallization just before the eruption. The single oldest grain could be either xenocryst or antecryst. Sakhno et al. (2010) report the presence of two zircon populations of similar U–Pb age (88.6 ± 1.3 and 85.80 ± 0.96 Ma) for the sample of a dacitic ignimbrite collected ca. 40 km northeast from the T05-66 sample (Fig. 2). In addition, three samples (04–418, 2238a/02 and 2320a/02) yield a relatively high MSWD of 4.1 to 6.4; this could also result from the involvement of minor amounts of the inherited Pb, having the antecrystic or xenocrystic nature. The age variations of individual zircons from these three samples do not exceed the analytical errors (Table 1), suggestive the effect of an inherited component to be insignificant. For the whole isotopic age dataset presented on Fig. 4, the only internal discrepancy was detected for the sample from the Pykarvaam suite (north CCS area, Fig. 4a) reported by Kelley et al. (1999). This sample reveals the 40Ar/ 39Ar age of 88.9±0.9 Ma, whereas the stratigraphically lower Kytapkai suite yields a set of 40Ar/ 39Ar ages of 87.08±0.21 to 87.59±0.22 Ma (Ispolatov et al., 2004). This inconsistency is likely related to analytical technique. The age of the Pykarvaam suite was determined by total fusion of biotite (rather than by step heating) and the age may be affected by some inherited argon. The discrepancy looks insignificant, and does not disturb the overall pattern of age data. 8.1. Volcanic episodes and spatial migration of OCVB activity New zircon U–Pb dates on the northern OCVB are summarized on a histogram (Fig. 8a), together with the published U–Pb and 40Ar/ 39Ar dates. For most samples, the isotopic ages range between 106 Ma and 79 Ma. Cumulative probability curve displays three major peaks at ca. 87.5 Ma, 84 Ma, and 80 Ma. For several well studied locations, considerable (up to 20 m.y.) lull periods in the volcanic activity have been revealed. These temporal gaps are manifested by substantially different isotopic ages of adjacent stratigraphic units (Figs. 2, 4a, b and e). The East Chukotka segment appears to be younger than ANS and CCS. The unit traditionally considered to represent its ‘lower

andesites’ yields the age of 88.1 ± 1.2 Ma, which is even younger than its lowermost silicic units (92.0 ± 2.0 Ma and 91.5 ± 1.6 Ma). Here, the oldest isotopic age obtained for OCVB rocks is 93.3 ± 0.2 Ma (Calvert, 1999). Finally, the only age of 67 ± 1 Ma has been obtained for a pantellerite from the ‘upper basaltic’ formation of the East Chukotka segment (Sakhno et al., 2010). Therefore, the existing data imply this segment of the OCVB both began and ceased its activity about 12 m.y. later than CCS or ANS. This contradicts all existing stratigraphic models (e.g., Belyi, 1994; Kotlyar and Rusakova, 2004; Varlamova et al., 2004) that suggest nearly synchronous formation of all OCVB segments. Also, there is an evident compositional difference between volcanic sequences of the 89–87 Ma age: both CCS and ANS experienced peak of silicic volcanism, whereas volcanism in the East Chukotka segment of the OCVB was largely andesitic. The specific timing of the East Chukotka segment could be related with its abnormal spatial direction (Fig. 1). The paleo-tectonic models for the northern Pacific (e.g., Cox et al., 1989) suggest the oceanic plates were moving westward during 115–100 Ma span, and towards NE after 85 Ma. Hence, the SE-directed plate boundary at the East Chukotka segment could have a transform nature at the early stages of the OCVB activity (106–95 Ma), and lack notable manifestations of volcanism, alike the present-day western part of the Aleutian arc (Rosenbaum and Mo, 2010). Volcanic activity here started later, when the normal component of plate convergence has become significant. Combination of radiometric dates and volcanic stratigraphy allows distinguishing five major stages in the history of the northern OCVB: (1) ca. 106–98 Ma; (2) ca. 94–91 Ma; (3) ca. 89–87 Ma; (4) ca. 85.5– 84 Ma; and (5) ca. 82–79(?) Ma. For each stage, the temporal limits have been determined using the cumulative probability plot (Fig. 8a). Our field observations, available geologic maps (Varlamova et al., 2004), and interpretation of satellite and aerial images allow outlining spatial extent of the four earlier episodes, which produced the volcanic pile of CCS and northern ANS (Fig. 2). During the first stage (106–98 Ma), the ‘lower andesites’ (over 2 km) were accumulated in the ANS. At the same time, a dominantly felsic sequence (the lower part of the Pucheveyem suite, up to 1.5 km thick) was formed in the western CCS, superposed on the extinct Tytylveyem volcanic belt (Fig. 2). Within the northern CCS, shoshonites and latites of the Etchikun suite (ca. 0.5 km) were erupted. Hence, the activity of the northern OCVB began in Albian (ca. 106 Ma),

24

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P.L. Tikhomirov et al. / Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32

Fig. 8. Histograms and cumulative probability plots for all recently published isotopic ages of both volcanic and plutonic rocks of the OCVB (excluding the results of K–Ar and bulk rock determinations which are not reliable enough). A) northern OCVB (area shown on Fig. 2), data after Moll-Stalcup et al., 1995; Kelley et al., 1999; Ispolatov et al., 2004; Tikhomirov et al., 2006; Sakhno et al., 2010; and this study; B) southern OCVB, data after Akinin and Miller, 2011, and references therein; C) entire OCVB, data from Calvert, 1999; Akinin and Calvert, 2002; and sources mentioned above.

both in CCS and ANS, with the strong compositional variations along the strike of the belt. The second episode (94–91 Ma) produced poorly stratified felsic sequences (over 1 km) at the base of the volcanic pile in the western part of the East Chukotka segment of the OCVB (Fig. 4g). Within the western CCS, the accumulation of the silicic Pucheveyem suite was continued, after the nearly 9 m.y. hiatus. In the northern ANS, the Upper Eropol suite (ca. 1 km) of rhyolitic tuffs and lavas with minor andesites and basalts was formed. Within the northern CCS, no products of this volcanic episode were found. Here, the Etchikun suite is directly overlain by volcanic strata related with the next stage of the OCVB activity. The third (89–87 Ma) and the fourth (85.5–84 Ma) pulses produced compositionally similar volcanic formations, which occupy about 70% of the area of CCS and northern ANS (Fig. 2). The duration of the quiescence between the two volcanic pulses (1.5 m.y.) is comparable with the analytical error of age determinations. Consequently, the existence of two separate pulses may be challenged. However, the two statistical peaks do exist, and the products of the third and the

25

fourth volcanic flare-ups appear separated spatially (Fig. 2). Most of samples with the isotopic ages of 89–87 Ma were collected in the northern part of the CCS (in the rear zone of the OCVB), whereas dates of 85.5–84 Ma characterize the southern CCS and the frontal zone of the OCVB. The plausible explanation is the extinction of volcanic activity in the northern CCS after 87 Ma. Within the ANS and southern CCS, two U–Pb dates of 88.2 ± 1.5 and 88.6 ± 1.3 Ma (Fig. 2) infer the presence of products of the 89–87 Ma stage, concealed by younger strata; this is consistent with the age of zircon antecrysts from the sample T05-66 (89.78 ± 0.77 Ma). The products of the fifth volcanic episode (82–79(?) Ma) include some local fields of silicic rocks within the Central Chukotka and East Chukotka segments, yet not well defined because of insufficient data. Among 30 isotopic dates available for CCS and northern ANS, only one corresponds to this time span (Fig. 2), and the volume of volcanic rocks accumulated during the fifth episode seems to be insignificant compared with that of previous four episodes. However, granodiorites of the Tanyurer batholith of East Chukotka segment (Fig. 2) which display 40Ar/ 39Ar ages of 79.7 ± 0.4 Ma (biotite) and 79.2 ± 1.9 Ma (amphibole) could have crystallized during this stage (Tikhomirov et al., 2006). Also, one may admit that some of 'upper basalts' of CCS and ANS could be related with the fifth episode, but any direct constraints on their age have not yet been obtained. In the southern part of the OCVB (Okhotsk segment), the 'upper basaltic' formations reveal the 40Ar/ 39Ar and zircon U–Pb ages between 79 and 74 Ma (Hourigan and Akinin, 2004; Akinin et al., 2007), and the only U–Pb date for the East Chukotka segment is as young as 67 Ma (see above). Taking into account evidence for a somewhat younger age of the East Chukotka segment, we suggest the ‘upper basalts’ of ANS and CCS were likely contemporaneous with those of the southern OCVB. Fig. 8b illustrates variations of published zircon U–Pb and 40Ar/ 39 Ar ages of magmatic rocks from the southern OCVB (Okhotsk and West Okhotsk segments). The ‘upper basalts’ display a distinct peak at 78–74 Ma (see above). For all older volcanic formations, the total timing seems to be quite similar to that of the northern OCVB. The oldest date is 106.2 ± 1.8 Ma, and the youngest sample from below the ‘upper basalts’ yields the age of 80.4 ± 2.1 Ma. Most dated samples return ages between 88 and 80 Ma, which correspond to two main volcanic episodes of the northern OCVB (89–87 and 85.5–84 Ma). That is, the period of the highest volcanic activity seems to be nearly synchronous throughout the whole OCVB. However, the statistical peaks on the histogram for the southern OCVB look less clear than those for the northern part of the belt. Also, the time span for the most prominent peaks is somewhat different (85–81 Ma at the south, and 89–87 Ma at the north). With the present-day data, it would be premature to decide, whether this results from any real variations of volcanic productivity, or from the paucity of sampling. The age histogram for the entire OCVB is presented on Fig. 8c. The dataset comprises 98 recently published zircon U–Pb and mineral separate 40Ar/ 39Ar ages, including results on several samples from the eastern part of the East Chukotka segment (Calvert, 1999; Akinin and Calvert, 2002). According to this histogram, volcanic complexes of 106–90 Ma age are relatively poorly exposed at the surface and reveal a statistical peak at ca. 93 Ma. More abundant are magmatic rocks with ages between 89 and 79 Ma; they take about 70% of the total set. Within this time interval, the cumulative curve displays peaks at ca. 87, 84, 82 and 80 Ma. Five samples younger than 79 Ma represent the ‘upper basaltic’ units. The estimated OCVB's lifetime totals to ca. 27 Ma (not including the ‘upper basalts’), or 32 Ma (including the ‘upper basalts’). Assuming that the anomalously young pantellerite of ca. 67 Ma (Sakhno et al., 2010) also belongs to the OCVB would

Fig. 7. a-n: Tera-Wasserburg concordia diagrams for the analysed zircons from volcanic rocks of the northern OCVB. Error ellipses are shown at 2 sigma uncertainty. The analyses excluded from the are calculation (because of discordance or high U content) are shown in dashed ellipses; o — Tera-Wasserburg concordia diagrams for zircons from the sample T05-66, ellipses at 2 sigma uncertainty. The dashed ellipse corresponds to zircon grain of a likely xenocrystic nature, and dotted ellipses mark the analyses of zircons interpreted as antecrysts (see discussion in the text); p — weighted average plot for the sample T05-66 (error bars are 2 sigma). The analytical results are listed in Table 1.

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increase the total duration of OCVB activity up to ca. 39 Ma, however this needs to be confirmed by with more representative dataset. The earliest formations of the OCVB have the same age (ca. 106–100 Ma) with the Magadan and East Taigonos granitic batholiths (ca. 106–97 Ma, after Farrar, 1992; Luchitskaya et al., 2003) located beyond the OCVB (Fig. 1) and presumably related with the Uda–Murgal arc. Therefore, the onset of the OCVB activity took place before the final extinction of Uda–Murgal arc magmatism. The less likely alternative is that the southern Taigonos and Magadan plutons represent the roots of a hypothetical deeply eroded part of the frontal zone of the OCVB. The revealed multi-pulse volcanism patterns may be caused by various reasons. The fast increase of magma production could result from changes in the slab geometry (e.g., Haschke et al., 2006), or in the rate of plate convergence (e.g., Hughes and Mahood, 2008), or from the subduction of a spreading ridge (e.g., Brown, 1998), or from the lithospheric delamination (e.g., Zhang et al., 2008). We also admit that the episodicity at the scale of several m.y. may be related to internal dynamics of crustal magma chambers. However, the discussion of possible effects of various processes is out of scope of this paper; we leave this problem for another publication. 8.2. Volume of volcanic rocks and extrusion rates Kotlyar et al. (1981) have reported the total volume of OCVB of 1.043 ⁎ 10 6 km 3, but they do not present neither the input data (e.g., outcrop areas and thicknesses of volcanic units), nor the details of calculation approach, hence any verification of this volume estimate is difficult. Besides, this estimate includes the volume of some complexes of the Uda–Murgal arc, which were formerly considered to belong to the OCVB. By this reasons, the published data on the total volume of the OCVB need a revision. The outcrop area of various volcanic complexes may be easily measured using the existing geologic maps (e.g., Varlamova et al., 2004) and

ArcGIS software. For volumetric calculations, the crucial and the most ambiguous variable is the total thickness of a volcanic pile. Available publications report rough estimates of the average thickness of the OCVB's volcanic pile of 2–4 km (see Section 3). Minimal average thickness of ca. 1.5 km is necessary to explain the absence of outcrops of the OCVB base over areas of several hundred km across, with relative topographic heights of up to 1.2 km. The additional clues on average thickness of OCVB can be obtained from structural observations. As noted in Section 3, the volcanic belt comprises extensive subsidence-related monoclines (Fig. 5) with dip angles of 5° to 15° and strike extent 100–150 km. The width of a monocline and the mean dip angle yield the increment of the total thickness of a volcanic pile of up to 3 km (Fig. 9). Also, the formation of calderas may cause a significant local increase of the thickness of a volcanic pile. In well-studied provinces of silicic volcanism (e.g., Cenozoic volcanic fields of Western USA), the estimated depth of caldera subsidence usually exceeds 2 km (Lipman, 2007), and some caldera-fill sequences may be as thick as 5 km (John, 1995; Lipman, 2007). The presence of volcanic units of a similar thickness within the OCVB, which comprises much greater volume of silicic volcanic rocks, seems quite possible. Actually, contribution of intracaldera volcanics to the total volume of the OCVB is difficult to estimate because of their uneven spatial distribution. Many calderas are probably concealed by younger outflow volcanics, and some subsidence structures are virtually undetectable being bounded by faults without any considerable tilting of strata. Taking into account the structural data, we suggest variations of the total thickness of the northern OCVB shown on the Fig. 9. These estimates are largely based on the position and structural characteristics of the most extended monoclines, which likely correspond to margins of wide subsidence structures. For the area where calderas are especially abundant ('nested'), we assume the increase of the average thickness by approximately 2 km. Beyond this area, calderas also occur, but they appear to be spatially separated, and their contribution to the total thickness of volcanics was omitted. Fully realizing

Fig. 9. Variations of the estimated total thickness of the OCVB northern part. 1–5 — average thickness of volcanic rocks: 1–0.5 km, 2–1.5 km, 3–2.5 km, 4–3.5 km, 5–4 km; 6 — area with abundant (‘nested’) calderas; 7 — present-day boundaries of the OCVB, exposed (a) and overlain by Cenozoic sediments (b); 8 — isolines of the estimated total thickness of volcanic rocks; 9 — the boundary of the OCVB part used for volumetric calculations; 10 — the largest monoclines (dip angles of 5 to 20°), which probably correspond with margins of wide volcanic subsidence structures (see discussion in the text); hatches indicate the dip of strata; 11 — main faults.

P.L. Tikhomirov et al. / Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32

relatively high uncertainty of these thickness estimates, we consider them as a first approach for the volumetric calculations, with the paucity of more reliable data. Our estimates imply the maximal total thickness of the northern OCVB of 4 km, and the average thickness of nearly 2.5 km. The most conservative assessment based on the relative heights of the presentday topography (which is surely an underestimate) yields at least 1.5 km. Any suggestions about the maximal possible average thickness of the OCVB are almost purely speculative. We assume the estimate by Sidorov et al. (2009) (3.5 km for the rear zone of the OCVB) to be an upper limit for this parameter. For the frontal zone of the volcanic belt, Sidorov and colleagues report the even greater value of 7 km, but these estimates were performed on the complexes which likely belong to the Uda–Murgal arc (see Section 3), so we use the estimate for the rear zone alone. According with the assumed average thickness, the calculated volume of volcanic rocks within the study area depicted on Figs. 2 and 9 (near 400 km along the NE-directed portion of the belt) ranges between 144,000 and 336,000 km3, with the most likely value of 240,000 km3 (Table 2). We suggest the real volume of the studied part of the OCVB may approach the maximal estimate or even exceed it, because the eroded volcanic complexes have not been taken into account. This eroded volume must be considerable because (1) the thickness of the volcaniclastic forearc sediments of the OCVB measured in local sections approaches 3 km (Varlamova et al., 2004), in spite of their limited present-day outcrop area; (2) the Tanyurer granitic pluton which intrudes the volcanics crystallized at depth of 6 ± 3 km (Alin-hornblende barometer, Tikhomirov et al., 2009b). For the entire OCVB, the estimated volume should be rather impressive, provided that the study area includes only about 12% of the total

27

length and 25% of the crop area of the volcanic belt. With the most conservatively estimated average total thickness of the OCVB (1.5 km), its calculated volume is ca. 550,000 km 3, or approximately 380,000 km3 in dense rock equivalent. Taking into account the present-day crop area of main rock types within the OCVB, we suggest the portions of the extrusive felsic and mafic/intermediate rocks to be roughly equal. The assumption of the minimal plutonic/volcanic volume ratio of 3:1 for silicic rocks (de Silva and Gosnold, 2007) yields the overall volume of OCVB's crust-derived magmas of about 760,000 km3. The model of crustal melting induced by basaltic underplating (Huppert and Sparks, 1988; Bergantz, 1989) suggests that the generation of a volume unit of anatectic rhyolitic magma requires equal or greater volume of mantle-derived magmas ponded at the base of, or within the continental crust. Consequently, the total magmatic addition to the crust during the OCVB formation must exceed 1.7 × 10 6 km3. The new data allow semi-quantitative estimate of volcanic output rates for the observed part of the OCVB. Table 2 lists calculation parameters and compares results with similar data from well-studied silicic volcanic provinces of the Phanerozoic (after Bryan et al., 2008). The volcanics of the East Chukotka segment and the ‘upper basalts’ were excluded from calculations because of insufficient number of radiometric dates; Table 2 presents only volumetric parameters of the 'upper basalts'. The calculation approach was as follows. On the basis of available isotopic dates, geologic maps, and satellite images, the areas corresponding to different volcanic episodes were outlined (Fig. 2). We did not separate the third (89–87 Ma) and the fourth (85.5–84 Ma) stages, because volume proportions of their products are ambiguous. Volume of each volcanic episode was calculated from the total area of outcrops, and from the thicknesses presented on Fig. 9. The correction on the possible extent of volcanic sequences

Table 2 Extrusion rates calculated for the northern OCVB, in comparison with the data for some well-known silicic volcanic provinces of the Earth. Age (Ma) a

1st (the ‘lower 106–98 8 andesites’) 2nd (silicic) 94–91 3 89–83.5 5.5 3rd to th (dominantly silicic) the ‘upper 79–74 5(?) basalts’ (?) Total for the studied including the 106–74 32(?) part of the OCVB ‘upper basalts’ (?) w/o the 106– 22.5 ‘upper basalts’ 83.5 Volumetric and age data for other silicic volcanic provinces (after Whitsunday 132–95 37 (Eastern Australia) 188–153 35 Chon Aike (Patagonia and Antarctic Peninsula) Sierra Madre Occidental 38–20 18 (Mexico) Kennedy–Connors–Auburn 320–280 40 (NE Australia) Altiplano-Puna (Andes) 10–3 7 Taupo (New Zealand) 1.6–0 1.6

Volcanic episodes of the northern OCVB

Volume Average Area of the Length of the thickness (103 km3) volcanic area concealed part (assumed) (km) measured along (km2) the strike of the volcanic belt (km)

Average Average extrusion extrusion rate per 1 km rate (10− 5 km3//(yr*km)) (km3/k.y.)

12730

6370

4190 65910

2100 14240

7910

Duration of Crop area (km2) volcanic activity (m.y.)

1.6

31

190

3.8

2.0

1.4 2.5

9 200

200 400

2.9 36.0

1.5 9.0

0

0.6

5

120

–

–

90740

6330

2.5

245

400

–

–

90740

6330

1:5 2.5 3:5

240

400

10.7

2500

55

2.2

235

3000

7.1

0.2

390

2000

22

1.1

500

1900

12.5

0.7

d

1.4 3.7

a

144 336

6:4 15:0

2.6

1:6 3:6

Bryan et al., 2008) 2200

30 20

c

300 300

4.3 9.4 to 13

e

For the total volume and extrusion rate for the studied part of the OCVB, the most likely values are presented in a bold script. Also, the minimal (numerator) and the maximal (denominator) values are given, according with various estimates of the average thickness of the volcanic pile (1.5 km is the minimal value indicated by the present-day toporgaphy, and 3.5 km is the estimate after Sidorov et al., 2009). a For each stage, the time brackets have been determined using the cumulative probability plot (Fig. 8a). b Including the area overlain by the ‘upper basalts’ (7910 km2) and by Cenozoic clastic sediments (6330 km2). c This value includes 1.4 ⁎ 106 km3 of coeval volcanogenic sediments (Bryan et al., 2000) — the factor omitted for other provinces listed. d According to data presented by de Silva and Gosnold (2007), the average volcanic output rate is only 1 ⁎ 103 km3/k.y., but increases up to 12*103 km3/k.y. during volcanic flare-ups. e According to data presented by de Silva and Gosnold (2007), this value is about 0.3 ⁎ 10− 5 km3/(yr ⁎ km).

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under the overlying strata was made, implying the additional 50 vol.% for the first and the second stages, and 22% for the combined third and fourth stages. The average volcanic output rate calculated for the studied part of the OCVB (400 km-long segment of the 3200 km long belt) ranges between 6.4 and 15.0 km3/k.y., depending on the assumed total thickness of a volcanic pile; the optimal value is 10.5 km3/k.y. (Table 2). Per 1 km of the strike of the belt, the output rate is 1.6 to 3.6× 10− 5 km3yr− 1 km− 1, with the optimal value of 2.6 × 10− 5 km3yr− 1 km− 1. Even the lowest of possible values are comparable to the extrusion rates of largest silicic volcanic provinces of the Earth (Table 2). The assumption of the minimal average thickness of the OCVB (1.5 km) yields the total volcanic output rate of ca. 22 km 3/k.y., or 15 km 3/k.y. in dense rock equivalent. This implies the rate of magmatic addition to the crust of >70 km3/k.y., or >2.2 × 10− 5 km3yr− 1 km− 1. Actually, these values are likely underestimates. The structural data available for the studied northern part of the volcanic belt (Fig. 9) suggest the average thickness of volcanics substantially greater than 1.5 km. Other parts of the OCVB do not differ greatly from CCS and ANS by structure and incision level (except the relatively eroded Okhotsk segment), so they probably have a similar total thickness. The real present-day volume of the volcanic belt may exceed 106 km3, as estimated by Belyi (1977) and Kotlyar et al. (1981), though we admit all these estimates are quite imprecise. To put further constraints on the original volume of the OCVB, it is pertinent to take into account the volcanic sequences recently concealed by Okhotsk and Bering seas (Fig. 1), and the eroded volcanic complexes. At a first approximation, this additional volume may be comparable with that of the remaining volcanic pile. Thus, the order-of-magnitude estimate for the original extrusive volume of the OCVB is ca. 2 × 10 6 km3, or about 1.4 × 106 km3 in dense rock equivalent. Assuming the crustal anatectic magmas take nearly 50% of this volume (see above), the total magmatic addition to the crust related with the OCVB formation could exceed 6 × 10 6 km3, which implies the magma production rate of 240 km 3/k.y., or 7.5 × 10− 5 km3yr− 1 km− 1. Such an event would require a deep thermal reworking of the crust. This suggestion is supported by U–Pb ages of zircons from middle and lower crustal xenoliths from the OCVB area (Akinin et al., 2009). All the 125 studied grains reveal ages between 138 and 60 Ma (i.e., presence of any pre-Cretaceous zircons have not been detected), and the most abundant are zircons of 100 to 80 Ma age. Among other silicic volcanic provinces, the only Whitsunday province of Eastern Australia has been reported to have an extrusion rate higher than that of the OCVB (55 km 3/k.y., Table 2). However, the calculated volcanic volume of this province (2.2 × 10 6 km 3) includes 1.4 × 10 6 km 3 of coeval volcaniclastic sediments from the neighbouring basin (Bryan et al., 2000). The comparison on an equal basis (e.g., the present-day crop area) indicates the OCVB is definitely larger than the Whitsunday province, and probably the largest province of silicic volcanism on the Earth (at least for the Phanerozoic). The Taupo zone known as the most active recent area of silicic volcanism has the calculated extrusion rate of 3.7 × 10− 5 km3yr− 1 km − 1 (Houghton et al., 1995), which exceeds the appropriate estimates for the OCVB. But the total volcanic volume of the Taupo zone and the duration of its activity are only about 2% and 7% of those of the Okhotsk–Chukotka belt, respectively (Table 2). During peak periods of OCVB's activity, the extrusion rate was likely much higher than that of Taupo zone, and any other present-day area of silicic volcanism. The value of 9.0 × 10 − 5 km 3yr − 1 km − 1 has been obtained for the combined third and fourth volcanic stages of the northern OCVB (Table 2); during each of these stages, the volcanic productivity could be substantially higher. 8.3. OCVB and silicic LIPs: a semantic discussion Silicic LIPs, or SLIPs, have been specified as a subgroup of LIPs characterized by high proportions of felsic volcanic rocks (over

75 vol.%, according to Bryan et al., 2002). By definition of a LIP by Bryan and Ernst (2008), all provinces of this kind must have a crop area >105 km2, total volume >105 km3, maximum lifespan of 50 m.y., intraplate tectonic setting or geochemical affinity, and highly episodic activity. For the Phanerozoic, only four SLIPs have been reported: Whitsunday and Kennedy–Connors–Auburn provinces of Eastern Australia, Chon Aike province of Patagonia and Antarctic Peninsula, and Sierra Madre Occidental province of Mexico (Bryan et al., 2002). The OCVB, with its total area of ca. 4.5 × 10 5 km 2, extrusive volume of over 10 6 km 3 (see Section 8.2), the lifetime of about 30 m.y., and with the recently detected volcanic flare-ups in its history fits both temporal and volumetric criteria of a LIP. However, the portion of silicic rocks in volcanic formations of the OCVB is below 75%, though some major segments of the volcanic belt (e.g., CCS) may contain up to 80–85% of dacites and rhyolites. Our estimates using the geologic map of 1:1,500,000 scale (Geologic map of the Northeast of USSR…, 1980) yield the portion of silicic volcanics of nearly 55% of the entire outcrop area of the OCVB. The rest 45% are dominated by mafic and intermediate rocks of typical calc-alkaline series (Polin and Moll-Stalcup, 1999; Kalinina et al., 2008; Akinin and Miller, 2011). Therefore, OCVB seems to represent a transition between the 'recognized' SLIPs and continental volcanic arcs, which may also comprise voluminous silicic formations (e.g., de Silva, 1989). Here we face the contradiction. Both the traditional understanding of LIPs and their revised definition proposed by Bryan and Ernst (2008) imply their intraplate setting, or characteristic intraplate geochemical affinity; but the OCVB, having all temporal and volumetric characteristics of a LIP, is a continental volcanic arc. This volcanic belt has been formed in a convergent margin zone (e.g., Parfenov, 1991; Nokleberg et al., 2001), and it is accompanied by remnants of related forearc basin and accretionary complexes (Varlamova et al., 2004; Soloviev et al., 2006). Finally, most mafic and intermediate formations of the OCVB have clear geochemical features of subduction-related magmatic complexes. The above-mentioned data lead to further discussion of the term ‘Large Igneous Province’ (e.g., Coffin and Eldholm, 1992; Sheth, 2007; Bryan and Ernst, 2008). It would be pertinent to note that the OCVB is not a unique large subduction-related volcanic province having a relatively high portion of silicic rocks. There are several other Phanerozoic volcanic belts having similar volumetric, temporal, and compositional characteristics, e.g., Kazakhstan and Balkhash-Ili belts, SE China belt, and East Sikhote–Alin belt. All these provinces are considered to be continental volcanic arcs (Tikhomirov, 2010, and references therein). Moreover, the geologic position of the ‘recognized’ SLIPs is not so much different from that of the continental subduction-related provinces, because all four ‘recognized’ Phanerozoic SLIPs have been formed relatively close (several hundred km) from active subduction zones. Such spatial proximity as well as calc-alkaline affinity of mafic rocks from typical SLIPs (Bryan et al., 2002), allow a possibility of genetic link (possibly indirect) with convergent plate margins. Thus, the magmatic systems of active continental margins may produce volcanic provinces with size and output rate similar to those of typical LIPs. Considering the intraplate setting as an obligatory attribute of LIPs, we need to classify such subduction-related provinces as ‘another LIPs’. However, if the volume and eruption rate are the main criteria, it has to be admitted that some LIPs may be genetically related to subduction environments. 8.4. Origin of the OCVB and implications on the genesis of large provinces of silicic volcanism Possible reasons for remarkably high volumes of silicic rocks in volcanic formations of the OCVB and similar subduction-related belts need to be addressed. Conventionally, the voluminous silicic volcanism is thought to result from melting of continental crust, with the underplated mantle-derived mafic magmas being the main heat source. The events of this kind may take place at active margins and in various

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continental intraplate environments (e.g., de Silva, 1989; Best and Christiansen, 1991; Riley et al., 2001; Ferrari et al., 2002; Lipman, 2007; Bryan and Ernst, 2008). Among the prerequisites for the formation of large provinces of silicic volcanism discussed in publications (e.g., Hildreth, 1981; Thompson, 1996; Bryan et al., 2002; Hughes and Mahood, 2008; de Silva, 2008), three seem to be the most popular: (1) extensional tectonic setting (which may be related to a continental break-up, or have a backarc nature); (2) composition of crustal protolith (hydrated amphibolitic lower crust is able to produce greater amounts of silicic melts than ‘dry’ granulitic protolith, under the same conditions); (3) an elevated output of mantle-derived magmas. In case of OCVB role of extension is unlikely to be significant. Structural studies do not identify evidence for any pronounced extensional regime during OCVB activity (Miller et al., 2002), and geophysical data do not reveal any crustal thinning under the volcanic belt (Nikolaevsky, 1967; Simonov et al., 2003). Considering the structural features of other large silicic volcanic provinces, one may conclude that considerable syn-volcanic extension is not a necessary condition for their formation. Despite some provinces do have some tracks of such a tectonic background (e.g., Wilson et al., 1995; Ferrari et al., 2002), others do not (Best, Christiansen, 1991). The effect of water content in crustal protolith could have been significant during the OCVB formation, because this belt is largely superposed on complexes of Paleozoic to Early Mesozoic active margins. At lower crustal levels, these complexes are expected to be substantially amphibolitic, not granulitic. Thus hydrated composition of crustal protolith may have contributed to strongly silicic composition of OCVB and similar provinces (Stephens et al., 1995; Bryan et al., 2002). Elevated flux of basaltic magma could also be responsible for the generation of huge volume of crust-derived magmas of the OCVB. Nevertheless, the estimated total magma production rate for the OCVB (7.5 × 10 − 5 km 3yr − 1 km − 1) does not exceed that for some island arcs which do not seem to be extra active – for instance, the Aleutian arc (9 to 18× 10− 5 km 3yr− 1 km− 1; Jicha et al., 2006). This leaves the possibility that a province like the OCVB may be formed

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at a subduction zone characterized by more or less common magma production rate. If so, other agents should be considered to explain the observed high output of silicic magmas. In this respect, we would like to focus on the thermal state of the crust, one of crucial parameters which control the crustal melting (Brown, 2001; Chen and Grapes, 2007, and references therein). The link between the degree of underplating-induced melting of crustal protolith and the ambient temperature before the emplacement of mafic magmas is easily understandable, and was supported by numerical modelling (e.g., Bergantz, 1989; Annen et al., 2006). For example, the latent heat of melting of the lower crust material by 20% is about 70 kJ/kg, whereas the increase of its temperature by 100 °C requires 139 kJ/kg (data from Annen et al., 2006). That is, the heating of the protolith to the melting point consumes much more energy than the melting process itself. By this reason, the duration of a magmatic pause between a volcanic event and a previous major episode of crustal magmatism may be an important factor affecting crustal anatexis. The heat energy of previous events may significantly accelerate the melting process and increase the volume of silicic magmas. Correlation between duration of a magmatic lull before major volcanic event and the proportion of silicic rocks in corresponding volcanic section is illustrated by Fig. 10. The magmatic pauses which predate the formation of SLIPs are relatively short, 0 to ca. 100 m.y., and below 20 m.y. for most provinces. For the OCVB and other substantially silicic subduction-related volcanic belts listed in the previous section, duration of these pauses does not exceed 15 m.y. On the contrary, most essentially mafic LIPs (Karoo, Deccan, Siberia, Parana-Etendeka etc.) display much longer magmatic gaps of 200 to 650 m.y. On the scale of a single volcanic province, spatial compositional variations sometimes display similar relations. For example, the average SiO2 content in volcanic rocks strongly increases from ANS to the adjacent CCS (Figs. 3 and 4). This observation corresponds with the records of major magmatic events in the CCS area just before the inception of OCVB activity, at 121–112 Ma (Tikhomirov et al., 2009a) and 107–105 Ma (Tikhomirov et al., 2011). In a similar way, the

Fig. 10. The age of the most known post-Paleozoic LIPs, compared with the age of a last previous major magmatic event in the same area (modified after Bryan and Ernst, 2008). Data sources are listed in Bryan and Ernst (2008) and Tikhomirov (2010). For the Balkhash-Ili and Southeast China volcanic belts, the age of main volcanic pulses is shown.

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basement of the northeastern (the most silicic) segment of the Devonian Kazakhstan belt includes large Late Silurian granitic plutons, whereas other parts of this belt overlie magmatic complexes of Ordovician or older age (Bakhteev, 1987). The effect of a ‘thermal and mechanical maturation of the crust’ as a factor promoting large-volume silicic eruptions has already been mentioned in publications (Lipman, 2007; de Silva, 2008). However, it has not been stated to be among the main factors of the formation of SLIPs and similar provinces. The remarkable volumetric and compositional characteristics of such provinces, along with their metallogenic significance and possible (perhaps, still underestimated) effect on the atmosphere and climate imply their origins need a more thorough understanding, with a more detailed view on the contribution from various genetic factors. 9. Summary 1. The OCVB is a continental volcanic arc, which has a remarkably high volume proportion (nearly 50%) of silicic rocks. Its volumetric and temporal parameters fit those of a LIP (Bryan and Ernst, 2008). Compositionally, the OCVB represents a ‘missing link’ between silicic LIPs and typical subduction-related continental arcs. 2. Generalized section of the OCVB comprises 3 main parts: (1) ‘lower andesites’, (2) various units dominated by silicic rocks, and (3) ‘upper basalts’. The present-day topography allows minimal estimate of the average thickness of a volcanic pile of about 1.5 km. For the area of a detailed study (400 km segment of the 3200 km long belt), the structural data yield the average thickness of ca. 2.5 km. 3. Combination of published and new U–Pb and 40Ar/ 39Ar age data indicates the OCVB was active since 106 Ma till ca. 79 Ma, excluding the ‘upper basalts’ that are believed to have been superposed on the OCVB. For the volcanic belt as a whole these ages agree with previously existed geochronological models (e.g., Filatova, 1988; Belyi, 1994) however stratigraphic models of some parts of the volcanic belt require major revisions. The most intense volcanic activity took place between 89 and 80 Ma (Coniacian through Early Campanian time), likely throughout the entire belt. This time span accounts for about 80% of total volume of the studied northern part of the OCVB. 4. The estimated extrusion rate for the northern OCVB strongly depends on the assumed average thickness of the volcanic pile, and ranges from 1.6 to 3.6 × 10 − 5 km 3yr − 1 km − 1, with the optimal value of 2.6 × 10 − 5 km 3yr − 1 km − 1. These values are comparable to those of SLIPs and most active recent subduction-related provinces of the Earth. 5. The OCVB activity was discontinuous. Most prominent activity peaks correspond to epochs of silicic volcanism, which display the volcanic output rate of up to 9.0 × 10 − 5 km 3yr − 1 km − 1, or even higher. 6. Proportion of silicic rocks in the volcanic section increases where basement of the volcanic area retained thermal energy of previous magmatic events. This factor may account for the observed spatial compositional variations of volcanic rocks in continental provinces, and may serve as an important genetic prerequisite for SLIPs in general. Acknowledgements Authors are grateful to Mr. J.-P. Londero (Kinross Gold, Toronto, Canada) for his friendly help with the organization of fieldwork, and to V.O. Ispolatov (Barrick Gold of Australia Ltd.) who kindly improved English and gave valuable comments on the manuscript. P. Tikhomirov thanks R. Tanaka, T. Ota and C. Sakaguchi (Pheasant Memorial Laboratory, Misasa, Japan) for their assistance with the preparation of samples, and all members of the Pheasant Memorial Laboratory for

the friendly and constructive environment. This study was funded by the Ministry of Education, Sports, Science and Technology of Japan, the program “Center of Excellence for the 21st Century in Japan” (E.N.) from the Japanese Society for the Promotion of Science, and in part by grants from the Ministry of Education, Science, Sports and Culture of Japan (T.M.). The work of P.T. was partially supported by Russian Foundation for Basic Research, grants 03-05-64623-a and 09-05-01197-a.

References Akinin, V.V., Calvert, A.T., 2002. Cretaceous mid-crustal metamorphism and exhumation of the Koolen gneiss dome, Chukotka Peninsula, northeastern Russia. In: Miller, E.L., Grantz, A., Klemperer, S.L. (Eds.), Tectonic evolution of the Bering Shelf – Chukchi Sea – Arctic Margin and adjacent landmasses: Geological Society of America Special Paper, 360, pp. 147–165. Boulder, Colorado. Akinin, V.V., Miller, E.L., 2011. Evolution of calc-alkaline magmas of the Okhotsk– Chukotka volcanic belt. Petrology 19 (2), 1–42. Akinin, V.V., Al'shevsky, V.V., Parfenov, M.I., Leonova, V.V., Conference proceedings dedicated to the memory of academician K.V.Simakov., 2007. Age and evolution of basalt magma from Ola plato, Magadan area. In: Chereshnev, I.A. (Ed.), Conference proceedings dedicated to the memory of academician K.V.Simakov., 5. SVKNII DVO RAN, Magadan (in Russian). Akinin, V.V., Miller, E.L., Wooden, J., 2009. Petrology and Geochronology of Crustal Xenoliths from the Bering Strait Region: Linking Deep and Shallow Processes in Extending Continental Crust. Geological Society of America Special Paper 456, 39–68. Annen, C., Blundy, J.D., Sparks, R.S.J., 2006. The genesis of intermediate and silicic magmas in deep crustal hot zones. Journal of Petrology 47, 505–539. Bakhteev, M.K., 1987. Paleozoic Orogenic Volcano-Plutonic Belts. Nedra, Moscow. 168 pp. (in Russian). Baranov, M.A., 1996. Nappe tectonics of the Myrgovaam Basin in Northwestern Central Chukotka. Russian Journal of Pacific Geology 15, 441–448 (In Russian, abstr. Eng.). Belyi, V.F., 1977. Stratigraphy and Structures of the Okhotsk–Chukotka Volcanic Belt. Nauka, Moscow. 190 pp. (in Russian). Belyi, V.F., 1994. The Geology of Okhotsk–Chukotka Volcanic Belt. SVKNII DVO RAN, Magadan. 76 pp. (in Russian). Belyi, V.F., Yefimova, A.F., Paraketsov, K.V., 1966. Lower Cretaceous of Okhotsk– Chukotka volcanic belt. International Geology Review 8 (10), 1226–1236. Bergantz, G.W., 1989. Underplating and partial melting: implications for melt generation and extraction. Science 245, 1093–1095. Best, M.G., Christiansen, E.H., 1991. Limited extension during peak Tertiary volcanism, Great Basin of Nevada and Utah. Journal of Geophysical Research 96 (B8), 13509–13528. Brown, M., 1998. Ridge–trench interactions and high-T–low-P metamorphism, with particular reference to the Cretaceous evolution of the Japanese Islands. In: Treloar, P.J., O'Brien, P.J. (Eds.), What Drives Metamorphism and Metamorphic Reactions: Geological Society Special Publication, 138, pp. 131–163. Brown, M., 2001. Crustal melting and granitic magmatism: key issues. Physical Chemistry Earth 26 (4–5), 201–212. Bryan, S.E., Ernst, R.E., 2008. Revised definition of Large Igneous Provinces (LIPs). EarthScience Reviews 86, 175–202. Bryan, S.E., Ewart, A., Stephens, C.J., Parianos, J., Downes, P.J., 2000. The Whitsunday volcanic province, central Queensland, Australia: lithological and stratigraphic investigations of a silicic-dominated large igneous province. Journal of Volcanology and Geothermal Research 99, 55–78. Bryan, S.E., Riley, T.R., Jerram, D.A., Leat, P.T., Stephens, C.J., 2002. Silicic volcanism: an under-valued component of large igneous provinces and volcanic rifted margins. In: Menzies, M.A., Klemperer, S.L., Ebinger, C.J., Baker, J. (Eds.), Magmatic Rifted Margins: Geological Society of America Special Paper, 362, pp. 99–118. Bryan, S.E., Ferrari, L., Reiners, P.W., Allen, C.M., Petrone, C.M., Ramos-Rosique, A., Campbell, I.H., 2008. New insights into Crustal contributions to large-volume rhyolite generation in the Mid-Tertiary Sierra Madre occidental province, Mexico, revealed by U–Pb geochronology. Journal of Petrology 49, 47–77. Calvert, A.T., 1999. Metamorphism and exhumation of mid-crustal gneiss domes in the Arctic Alaska terrane: Ph.D. thesis. Santa Barbara: University of California. 198 pp. Cas, R.A.F., 1983. A review of the Palaeogeographic and tectonic development of the Palaeozoic Lachlan Fold Belt of Southeastern Australia. Geological Society of Australia Special Publication 10 104 pp. Charlier, B.L.A., Wilson, C.J.N., Lowenstern, J.B., Blake, S., Van Calsteren, P.W., Davidson, J.P., 2004. Magma generation at a large, hyperactive silicic volcano (Taupo, New Zealand) revealed by U–Th and U–Pb systematics in zircons. Journal of Petrology 46, 3–32. Chen, G.-N., Grapes, R.H., 2007. Granite genesis: in-situ melting and crustal evolution. Springer. 218 pp. Coffin, M.F., Eldholm, O., 1992. Volcanism and continental break-up: a global compilation of large igneous provinces. In: Storey, B.C., Alabaster, T., Pankhurst, R.J. (Eds.), Magmatism and the Causes of Continental Break-up: Geological Society of London Special Publication, 68, pp. 17–30. Cox, A., Debiche, M.G., Engebretson, D.C., 1989. Terrane trajectories and plate interactions along continental margins in the North Pacific basin. In: Ben-Avraham, I. (Ed.), The evolution of the Pacific Ocean margins. Oxford University Press.

P.L. Tikhomirov et al. / Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32 de Silva, S.L., 1989. The Altiplano–Puna volcanic complex of the Central Andes. Geology 17, 1102–1106. de Silva, S.L., 2008. Arc magmatism, calderas, and supervolcanoes. Geology 36, 671–672. de Silva, S.L., Gosnold, W.D., 2007. Episodic construction of batholiths: insights from the spatiotemporal development of an ignimbrite flare-up. Journal of Volcanology and Geothermal Research 167, 320–335. Farrar, E., 1992. Thermal reworking of granitic rocks of the Okhotsk–Chukotka volcanic belt during Paleogene: evidences from 40Ar/39Ar dating. Russian Journal of Pacific Geology 11 (1), 109–114 (in Russian, abstr. Eng.). Ferrari, L., Lopez-Martinez, M., Rosas-Elguera, J., 2002. Ignimbrite flare-up and deformation in the southern Sierra Madre Occidental, western Mexico: implications for the late subduction history of the Farallon plate. Tectonics 21 (4). doi:10.1029/2001TC001302. Filatova, N.I., 1988. Circum-oceanic volcanic belts. Nedra, Moscow. 264 pp. (in Russian). Geologic map of the Northeast of USSR (scale 1 : 1 500 000), 1980. Gorodinsky, M.E. (ed.) Leningrad, VSEGEI. Haschke, M., Günther, A., Melnik, D., Echtler, H., Reutter, K.-J., Scheuber, E., Oncken, O., 2006. Central and Southern Andean deformation inferred from arc magmatism. In: Oncken, et al. (Ed.), The Andes – Active Subduction Orogeny. Springer, pp. 337–353. Herman, A.B., 1999. Cretaceous Flora of the Anadyr–Koryak Subregion (Northeastern Russia): Taxonomy, Age, Stratigraphic and Florogenetic Significance. Transactions of the Geological Institute of the Russian Academy of Sciences, 529. GEOS, Moscow. 122 pp. (in Russian). Hildreth, E.W., 1981. Gradients in silicic magma chambers: implications for lithospheric magmatism. Journal of Geophysical Research 86, 10153–10192. Houghton, B.F., Wilson, C.J.N., McWilliams, M.O., Lanphere, N.A., Weaver, S.D., Briggs, R.M., Pringle, M.S., 1995. Chronology and dynamics of a large silicic magmatic system: Central Taupo Volcanic Zone, New Zealand. Geology 23, 13–16. Hourigan, J.K., Akinin, V.V., 2004. Tectonic and chronostratigraphic implications of new 40 Ar/39Ar geochronology and geochemistry of the Arman and Maltan-Ola volcanic fields, Okhotsk–Chukotka volcanic belt, northeastern Russia. Geological Society of America Bulletin 116 637–354. Hughes, G.R., Mahood, G.A., 2008. Tectonic controls on the nature of large silicic calderas in volcanic arcs. Geology 36, 627–630. Huppert, H.E., Sparks, R.S.J., 1988. The generation of granitic magmas by intrusion of basalt into continental crust. Journal of Petrology 29, 599–642. Ispolatov, V.O., Tikhomirov, P.L., Heizler, M., Cherepanova, I.Yu., 2004. New 40Ar/39Ar ages of Cretaceous continental volcanics from Central Chukotka: implications for initiation and duration of volcanism within the northern part of the Okhotsk– Chukotka volcanic belt (Northeastern Eurasia). Journal of Geology 112, 369–377. Jicha, B.R., Scholl, D.W., Singer, B.S., Yogodzinski, G.M., Kay, S.M., 2006. Revised age of Aleutian Island Arc formation implies high rate of magma production. Geology 34, 661–664. John, D.A., 1995. Tilted middle Tertiary ash-flow calderas and subjacent granitic plutons, southern Stillwater Range, Nevada: Cross sections of an Oligocene igneous center. GSA Bulletin 107, 180–200. Kalinina, E.A., Tikhomirov, P.L., Moriguti, T., Nakamura, E., 2008. The north of the Okhotsk– Chukotka volcanic belt (NE Russia): sources of mafic magmas. Origin, Evolution and Dynamics of the Earth. COE-21 International Symposium MISASA-III abstr, pp. 155–156. Kelley, S.R., Spicer, R.A., Herman, A.B., 1999. New 40Ar/39Ar dates for Cretaceous Chauna Group tephra, northeastern Russia, and their implications for the geologic history and floral evolution of the North Pacific region. Cretaceous Research 20, 97–106. Kotlyar, I.N., Rusakova, T.B., 2004. Cretaceous magmatism and mineral deposits of the Okhotsk–Chukotka region: geological and geochronological correlation. SVKNII DVO RAN, Magadan. 152 pp. (in Russian). Kotlyar, I.N., Belyi, V.F., Milov, A.P., 1981. Major element chemistry of igneous formations of the Okhotsk–Chukotka volcanic belt. Nauka, Moscow. 223 pp. (in Russian). Lawver, L.A., Grantz, A., Gahagan, L.M., 2002. Plate kinematic evolution of the present Arctic region since the Ordovician. In: Miller, E.L., Grantz, A., Klemperer, S.L. (Eds.), Tectonic Evolution of the Bering Shelf – Chukchi Sea – Arctic Margin and Adjacent Landmasses: Geological Society of America Special Papers, 360, pp. 333–358. Lipman, P.W., 1992. Magmatism in the Cordilleran United States; progress and problems. In: Burchfiel, B.C., Lipman, P.W., Zoback, M.L. (Eds.), The Cordilleran Orogen: Conterminous U.S. The Geology of North America. Geological Society of America, Boulder, pp. 481–514. Lipman, P.W., 2007. Incremental assembly and prolonged consolidation of Cordilleran magma chambers: evidence from the Southern Rocky Mountain volcanic field. Geosphere 3 (No. 1), 42–70. doi:10.1130/GES00061.1. Luchitskaya, M.V., Hourigan, J., Bondarenko, G.E., 2003. New SHRIMP U–Pb zircon data on granitoids from the Pribrezhnyi and Eastern Taigonos belt, southern Taigonos Peninsula. Doklady Earth Sciences 389A (3), 354–357. Ludwig, K.R., 2000. ISOPLOT. Version 2.3. Berkeley Geochronology Center, special publication, p. 1a. Miller, E.L., Verzhbitsky, V.E., 2009. Structural studies near Pevek, Russia: implications for formation of the East Siberian Shelf and Makarov Basin of the Arctic Ocean. Stephan Mueller Special Publication Series 4, 223–241. Miller, E.L., Gelman, M., Parfenov, L., Hourigan, J., 2002. Tectonic setting of Mesozoic magmatism: a comparison between northeastern Russia and the North American Cordillera. In: Miller, E.L., Grantz, A., Klemperer, S.L. (Eds.), Tectonic evolution of the Bering Shelf – Chukchi Sea – Arctic Margin and Adjacent Landmasses: Geological Society of America Special Paper, 360, pp. 313–332. Boulder, Colorado. Miller, E.L., Katkov, S.M., Strickland, A., Toro, J., Akinin, V.V., Dumitru, T.A., 2009. Geochronology and thermochronology of Cretaceous plutons and metamorphic country rocks, Anyui–Chukotka fold belt. North East Arctic Russia. Stephan Mueller Special Publication Series 4, 157–175. Mishin, L.F., Akinin, V.V., Mishin, E.L., 2008. New data on the age of magmatic rocks of the western segment of the Okhotsk–Chukotka volcanic belt. Russian Journal of Pacific Geology 27 (No. 5), 12–24 in Russian, abstr. Eng.

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Moll-Stalcup, E.J., Krogh, T.E., Kamo, S., Lane, L.S., Cecile, M.P., Gorodinsky, M.E., 1995. Geochemistry and U–Pb geochronology of arc-related magmatic rocks, Northeastern Russia. Geological Society of America Abstracts with Programs 27, 65. Montin, S.A., 1992. Folding and thrusting in Upper Cretaceous sequences of the Okhotsk–Chukotka Volcanic Belt. Vestnik MGU. Ser. 4. Earth Sciences No. 2, 88–91 (in Russian). Morozov, O.L., 2001. Geology and structure of Central Chukotka. Transactions of the Geological Institute of the Russian Academy of Sciences 523. GEOS, Moscow. 200 pp. (in Russian). Nikolaevsky, A.A., 1967. Deep structure of the Okhotsk–Chukotka volcanic belt. In: Gaisky, V.N., Karataev, G.I., Kaufman, A.A., Moiseenko, F.S., Puzyrev, N.N., Surkov, V.S. (Eds.), Regional geophysical research in Siberia. Novosibirsk, Nauka, pp. 137–148. in Russian. Nokleberg, W.J., Parfenov, L.M., Monger, J.W.H., Norton, I.O., Khanchuk, A.I., Stone, D.B., Scotese, C.R., Scoll, D.W., Fujita, K., 2001. Phanerozoic tectonic evolution of the Circum-North Pacific. US Geological Survey, Professional paper 1626. 122 pp. Obrutchev, A.V., 1934. Ore deposits. Gorgeonefteizdat, Moscow. 596 pp. (in Russian). Parfenov, L.M., 1991. Tectonics of the Verkhoyansk–Kolyma Mesozoides in the context of plate tectonics. Tectonophysics 199, 319–342. Polin, V.F., Moll-Stalcup, E.J., 1999. Petrological and geochemical criteria of the structural environment for formation of the Chukchi segment, the Okhotsk–Chukotka volcanic belt. Russian Journal of Pacific Geology 18 (No. 4), 29–47 in Russian, abstr. Eng. Riley, T.R., Leat, P.T., Pankhurst, R.J., Harris, C., 2001. Origins of large volume rhyolitic volcanism in the Antarctic Peninsula Patagonia by crustal melting. Journal of Petrology 42, 1043–1065. Rosenbaum, G., Mo, W., 2010. Tectonic and magmatic responses to the subduction of high bathymetric relief. Gondwana Research. doi:10.1016/j.gr.2010.10.007. Sakhno, V.G., Polin, V.F., Akinin, V.V., Sergeev, S.A., Alenicheva, A.A., Tikhomirov, P.L., Moll-Stalcup, E.J., 2010. The diachronous formation of the Enmyvaam and Amguema–Kanchalan volcanic fields in the Okhotsk–Chukotka Volcanic Belt (NE Russia): evidence from isotopic data. Doklady Earth Sciences 434 (Part 1), 1172–1178. Sheth, H., 2007. Large Igneous Provinces (LIPs): definition, recommended terminology, and a hierarchical classification. Earth Science Reviews 85, 117–124. Sidorov, A.A., Volkov, A.V., Belyi, V.F., Alekseev, V.Yu., Kolova, E.E., 2009. The gold–silver Okhotsk–Chukotka volcanic belt. Geology of Ore Deposits 51 (6), 441–455. Simonov, A.P., Berzin, R.G., Suleimanov, A.K., 2003. New data on the deep crustal structure of Northeastern Russia: the key profile 2-DV. Razvedka i Okhrana Nedr. (No. 5), 27–31 (In Russian). Sokolov, S.D., Bondarenko, G.Ye., Morozov, O.L., Luchitskaya, M.V., 2001. The tectonics of the transition zone between the Verkhoyansk–Chukotka and Koryak–Kamchatka folded areas. MOIP Bulletin: Earth Sciences 76 (No. 6), 24–37 in Russian. Sokolov, S.D., Bondarenko, G.Ye., Morozov, O.L., Shekhovtsov, V.A., Glotov, S.P., Ganelin, A.V., Kravchenko-Berezhnoy, I.R., 2002. South Anyui suture, northeast Arctic Russia: facts and problems. Geological Society of America, Special Paper 360, 209–224. Sokolov, S.D., Bondarenko, G.Ye., Khudoley, A.K., Morozov, O.L., Luchitskaya, M.V., Tuchkova, M.I., Layer, P.W., 2009. Tectonic reconstruction of Uda–Murgal arc and the Late Jurassic and Early Cretaceous convergent margin of Northeast Asia– Northwest Pacific. Stephan Mueller Special Publication Series 4, 273–288. Soloviev, A., Garver, J.I., Ledneva, G., 2006. Cretaceous accretionary complex related to Okhotsk–Chukotka Subduction, Omgon Range, Western Kamchatka, Russian Far East. Journal of Asian Earth Sciences 27, 437–453. Stephens, C.J., Ewart, A., Bryan, S., Schön, R.W., 1995. Rift-related, large-volume silicic volcanism associated with Lower Cretaceous continental breakup, eastern Australia. Boulder, Colorado, International Union of Geodesy and Geophysics 21 General Assembly, Abstracts, p. A443. Stone, D.B., Layer, P.W., Raikevich, M.I., 2009. Age and paleomagnetism of the Okhotsk– Chukotka Volcanic Belt (OCVB) near Lake El'gygytgyn, Chukotka, Russia. Stephan Mueller Special Publication Series 4, 243–260. Struzhkov, S.F., Konstantinov, M.M., 2005. Gold and Silver Metallogeny of the Okhotsk– Chukotka Volcanic Belt. Nauchnyi Mir, Moscow. 318 pp. (in Russian). Tatsumi, Y., Eggins, S., 1995. Subduction zone magmatism. Frontiers in Earth Sciences. Blackwell Science, Oxford. 211 pp. Thompson, A.B., 1996. Fertility of crustal rocks during anatexis. Transactions of the Royal Society of Edinburgh, Earth Science 87, 1–10. Tikhomirov, P.L., 1996. Ore-bearing volcanic structures of Mramorny ore cluster (Chukotka area). Izvestiya VUZov: Geologiya i Razvedka (N 6), 56–68 (In Russian). Tikhomirov, P.L., 2010. Largest Phanerozoic Provinces of Silicic Volcanism: Tectonic Position and Genetic Prerequisities. Moscow University Geology Bulletin 65 (3), 194–202. Tikhomirov, P.L., Akinin, V.V., Ispolatov, V.O., Alexander, P., Cherepanova, I.Yu., Zagoskin, V.V., 2006. The Okhotsk–Chukotka volcanic belt: age of its Northern Part according to new Ar–Ar and U–Pb geochronological data. Stratigraphy and Geological Correlation 14 (5), 524–537. Tikhomirov, P.L., Kalinina, E.A., Kobayashi, K., Nakamura, E., 2008. Late Mesozoic silicic magmatism of the North Chukotka area (NE Russia): age, magma sources, and geodynamic implications. Lithos 105, 329–346. Tikhomirov, P.L., Luchitskaya, M.V., Kravchenko-Berezhnoy, I.R., 2009a. Cretaceous granitoids of NE Asia: rock chemistry, composition of rock forming minerals, and conditions of formation. Stephan Mueller Special Publication Series 4, 289–311. Tikhomirov, P.L., Kalinina, E.A., Kobayashi, K., Nakamura, E., 2009b. Tytylveyem volcanoplutonic belt, the Early Cretaceous magmatic province of NE Asia. The Geology of Polar Regions of the Earth. Abstracts of the XLII conference on tectonics, vol. 2. GEOS, Moscow, pp. 239–241 (in Russian).

32

P.L. Tikhomirov et al. / Journal of Volcanology and Geothermal Research 221–222 (2012) 14–32

Tikhomirov, P.L., Luchitskaya, M.V., Shats, A.L., 2011. Age of granitoid plutons, North Chukotka: problem formulation and new SHRIMP U–Pb zircon dates. Doklady Earth Sciences 440, 1363–1366. Tilman, S.M., Bogdanov, N.A., 1992. Tectonic map of the Northeastern Asia, scale 1:5000000. State Geodesy and Cartography Committee, Moscow. Ueda, H., Miyashita, S., 2005. Tectonic accretion of a subducted intraoceanic remnant arc in Cretaceous Hokkaido, Japan, and implications for evolution of the Pacific northwest. Island Arc 14, 582–598. Umitbaev, R.B., 1986. Okhotsk-Chaun Metallogenic Province. Nauka, Moscow. 286 pp. (in Russian). Usui, T., Kobayashi, K., Nakamura, E., 2002. U–Pb isotope systematic of micro-zircon inclusions: implications for the age and origin of eclogite xenolith from the Colorado Plateau. Proceedings of Japanese Academy, Series B 78, 51–56. Varlamova, V.A., Malysheva, G.M., Vyatkin, B.V., Zvizda, T.V., Zhukov, V.A., Kovalenko, A.V., Kazinsky, V.A., 2004. Information report on unaccomplished works “The creation of digital geologic maps of 1:500000 scale for the Chukotka Autonomous

Okrug area” (The monitoring of regional geologic research works, scale 1:500000). FGUGP “Georegion”, Anadyr (in Russian). Wilson, C.J.N., Houghton, B.F., McWilliams, M.O., Lanphere, M.A., Weaver, S.D., Briggs, R.D., 1995. Volcanic and structural evolution of Taupo Volcanic Zone, New Zealand: a review. Journal of Volcanology and Geothermal Research 68, 1–28. Wilson, M., Neumann, E.-R., Davies, G.R., Timerman, M.J., Heeremans, E., Larsen, B.T., 2004. Permo-Carboniferous magmatism and rifting in Europe. Geological Society, London, Special Publications, 223. 503 pp. Yarmolyuk, V.V., 1973. Volcanic subsidence structures of the Okhotsk–Chukotka volcanic belt. Nauka, Novosibirsk. 103 pp. (in Russian). Zhang, J.-H., Ge, W.-C., Wu, F.-Y., Wilde, S.A., Yang, J.-H., Liu, X.-M., 2008. Large-scale Early Cretaceous volcanic events in the northern Great Xing'an Range, Northeastern China. Lithos 102, 138–157. Zhulanova, I.L., Rusakova, T.B., Kotlyar, I.N., 2007. Geochronology and geochronometry of endogenic events through Mesozoic in Northeastern Asia. Nauka, Moscow. 358 pp. (in Russian).