Boninites through time and space: petrogenesis and geodynamic settings Текст научной статьи по специальности «Науки о Земле и смежные экологические науки»

GEODYNAMICS & TECTONOPHYSICS

PUBLISHED BY THE INSTITUTE OF THE EARTH'S CRUST SIBERIAN BRANCH OF RUSSIAN ACADEMY OF SCIENCES

ISSN 2078-502X

2016 VOLUME 7 ISSUE 2 PAGES 143-172

http://dx.doi.org/10.5800/GT-2016-7-2-0202

BONINITES THROUGH TIME AND SPACE: PETROGENESIS AND GEODYNAMIC SETTINGS

A. A. Shchipansky

Geological Institute of RAS, Moscow, Russia

Abstract: The article provides an overview of boninitic magmatism occurrences in space and time and shows that the boninite rock series were generated through the entire geological history of the Earth. In modern environments, the genesis of boninites is related to intra-oceanic subduction initiation. Boninites are typical members of suprasubduc-tion zone ophiolite sequences in the Phanerozoic fold belts and also present in the early Precambrian greenstone belts. A comparative study on compositions of the early Precambrian and Phanerozoic boninites indicate their evolution through time due to gradual transition from the early thick-plate tectonics to the modern thin-plate tectonics. A link between subduction initiation and mantle-plume impingement at the oceanic lithosphere is discussed.

Key words: boninite; subduction initiation; mantle plume; secular changing in the Earth geodynamics

Recommended by E.V. Sklyarov

For citation: Shchipansky A.A. 2016. Boninites through time and space: petrogenesis and geodynamic settings. Geodynamics & Tectonophysics 7 (2), 143-172. doi:10.5800/GT-2016-7-2-0202.

БОНИНИТЫ ВО ВРЕМЕНИ И ПРОСТРАНСТВЕ: ПЕТРОГЕНЕЗИС И ГЕОДИНАМИЧЕСКИЕ ОБСТАНОВКИ ОБРАЗОВАНИЯ

А. А. Щипанский

Геологический институт РАН, Москва, Россия

Аннотация: Бониниты получили широкую известность благодаря глубоководным исследованиям преддуго-вых областей современных зон плитовой конвергенции Юго-Западной Пацифики. Однако они имеют широкое распространение и в офиолитах складчатых поясов, которые традиционно рассматриваются в качестве океанической коры геологического прошлого. Поскольку бониниты не известны в срединно-океанических хребтах, неизбежно возникает вопрос о природе офиолитов.

Общепринято, что под бонинитами понимаются вулканические породы, которые удовлетворяют следующим критическим параметрам составов (в пересчете на сухой остаток) - SiÛ2>52 вес. %; MgO>8 вес. % и TiÛ2<0.5 вес. % [Le Bas, 2000]. Их классификация основана на различиях в химических, а не минералогических составах, и принято различать две крупные группы бонинитов - высококальциевые и низкокальциевые [Crawford et al., 1989]. С бонинитами пространственно и генетически связаны примитивные островодужные низко-Ti лавы, что предопределило необходимость выделения обособленной магматической серии, известной как бонинитовая серия [Pearce, Robinson, 2010]. Собственно бониниты являются наиболее фракционированной ветвью серии, которая берет начало в пикритовых низко-Ti расплавах. Характер распределения спектра малых элементов бонинитов наглядно показывает необычайно высокую степень деплетации ман-

тийного источника при одновременных свидетельствах их надсубдукционного генезиса, например отрицательных аномалиях Nb(Ta) и Ti. Спектры малых элементов бонинитовой серии таковы, что, во-первых, исключается участие какого-либо вклада в их петрогенезис материала континентальной коры и, во-вторых, требуется плавление мантийного источника, более деплетированного по сравнению с лерцолитовой мантией, генерирующей расплавы MORB. В то же время геохимия пород бонинитовой серии демонстрирует их отчетливую связь с толеитами островных дуг - структур, в которых происходит формирование ювенильных порций континентальной коры.

В статье обобщены литературные данные по 36 объектам находок бонинитов в современных обстанов-ках, офиолитах и раннедокембрийских зеленокаменных поясах. Показано, что породы бонинитовой серии формировались на протяжении всей геологической истории Земли.

Петрологическая уникальность пород бонинитовой серии состоит в том, что для их генезиса требуется сочетание различных факторов, которое может реализовываться только в определенных, и очень ограниченных по месту локализации, геодинамических обстановках. Во-первых, происхождение источника бонинито-вых магм требует предварительного истощения верхнемантийного резервуара одним или несколькими эпизодами экстракции базальтовых расплавов; т.е. источником являлась гарцбургитовая мантия. Во-вторых, лавы бонинитовой серии характеризуются заметной обогащенностью крупноионными литофильными элементами и легкими редкоземельными элементами по сравнению с несовместимыми высокозарядными ионами. Такие их геохимические характеристики указывают на активность водного флюида, который должен был быть инфильтрирован в мантийный источник бонинитовых расплавов. Несмотря на неопределенности в экспериментальном моделировании расплавов бонинитовой серии, составы которых зависят от многих факторов, включающих степень деплетации мантии и флюидный режим плавления, существует ясность в том, что для их генерации требуются аномально высокие температуры и присутствие водосодержащего флюида в заметном количестве. На основе современной теории декомпрессионного плавления верхней мантии были проведены расчеты условий генерации первичных расплавов бонинитовой серии различного возраста, что позволило установить отчетливый эволюционный тренд их изменения. Показано, что раннедокем-брийские бонинитовые серии формировались при более высоких степенях плавления гарцбургитовой мантии (30-40 %), а формирование мантийных расплавных колонн происходило на существенно больших глубинах (3.5-4.0 ГПа), чем в фанерозойском эоне (2.5-3.0 ГПа).

Исследования современных проявлений бонинитового вулканизма демонстрируют, что они локализованы только в зонах интраокеанической плитовой конвергенции, и нет ни одного доказанного примера, свидетельствующего об иных геодинамических обстановках их формирования. Благодаря многочисленным находкам пород бонинитовой серии, в настоящее время стало очевидным, что большинство офиолитов мира маркируют формации не древних срединно-океанических хребтов, а палеозоны спрединга в надсубдукционных обстановках на границах океанических плит геологического прошлого. Понимание геодинамической обстановки формирования бонинитовых серий было связано с тем, что офиолиты супрасубдукционных зон связаны с начальными стадиями возникновения интраокеанических островных дуг. С физической точки зрения, главным условием для начала субдукции является возникновение гравитационной нестабильности в океанической литосфере, приводящей к ее полному расколу или коллапсу, а следовательно, к декомпрессионному плавлению верхней мантии и инициации погружения одной части плиты под другую. Это явление, как и генетическая связь бонинитов с офиолитами, легло в основание «правила инициации субдукции» (subduction initiation rule, SIR) [Whattam, Stern, 2011].

Теоретически, коллапс литосферы может произойти в двух случаях: 1) когда в соприкосновение приходят плиты с разными термальными характеристиками, например при трансформном совмещении плит разного возраста - древней, холодной, и молодой, горячей [Stern, 2004]; 2) когда место инициации субдукции определяется плотностными неоднородностями на границах нормальной океанической литосферы и утолщенной океанической литосферы плюмовой природы, т.е. океанических плато или трассеров воздействия горячих точек - асейсмических хребтов или симаунтов [Niu et al., 2003]. Хорошо известно, что подъем мантийного плюма приводит к ослаблению прочности литосферы и может вызвать раскол континентов. Но, помимо этого явления, внедрение плюма в литосферу существенно изменяет ее плотностные характеристики. Привнос в верхние горизонты мантии и океаническую литосферу расплавов из обогащенного глубинного источника должен приводить к рефертилизации ранее деплетированной мантии. По мере охлаждения такой процесс будет вести к уплотнению переработанной мантийным плюмом верхней мантии, а возникший в области переработки новый сегмент литосферы со временем может приобрести отрицательную плавучесть. Это обусловлено тем, что вулканиты OIB заметно обогащены Fe и Ti. Кроме того, хорошо известно, что Fe-Ti базальты/габбро эклогитизируются гораздо быстрее их магнезиальных эквивалентов.

По-видимому, процесс установления стационарного режима субдукции требует некоторого периода аккомодации, связанного с обрывами слэба и, как следствие, контрастностью тектонических режимов на поверхности. Причиной малоглубинного отрыва слэба могла стать плотностная неоднородность погружавшейся литосферы, например ее локальная переутяжеленность продуктами OIB магматизма. Важнейшими геодинамическими следствиями этого являются, во-первых, кратковременное сильное термальное возмущение над узколокализованной областью слэбового окна и, во-вторых, быстрый аплифт ее надсубдукционной области. Такой механизм хорошо объясняет кратковременность (3-5 млн лет) и большие объемы вулканизма, существенно превышающие объемы вулканизма в режимах стационарной субдукции [Stern, 2002, 2004]. Аплифт надсубдукционной области приводит к образованию на месте висячей плиты офиолитовой «платформы» - фундамента для островодужной постройки.

В раннем докембрии бонинитовый магматизм представлен широко, а количество новых находок древних бонинитов неуклонно возрастает. Согласно недавно опубликованным оценкам, объем бонинитового магматизма в архее примерно соответствует объемам коматиитов [Furnes et al., 2014]. Установление пород бонинитовой серии, ассоциирующих с фрагментами параллельных даек и метабазитами IAT-типа в древнейшем

сохранившемся комплексе Исуа, по-видимому, указывает на то, что процессы субдукции имеют корни, простирающиеся к началу геологической истории Земли. Поскольку процессы инициации субдукции требуют раскола океанической литосферы на ее полную мощность, раннедокембрийская литосфера по реологическим свойствам до ее основания должна была находиться в области хрупких или хрупко-пластических деформаций. Другими словами, такую литосферу можно рассматривать как жесткое тело, способное противостоять конвективной нестабильности, что является атрибутом плитовой тектоники [Sleep, 1992]. Мощность архейской океанической литосферы оценивается в 85-120 км, тогда как современной - примерно в 60 км.

В отличие от фанерозойских бонинитовых серий, родоначальные расплавы раннедокембрийских серий формировались на глубинах ~120-130 км, т.е. в поле стабильности алмаза. Учитывая то, что примитивные расплавы древних бонинитовых серий несут метки субдукционного влияния, можно думать о способности глубокого погружения слэбов в раннедокембрийскую мантию. Таким образом, можно полагать, что в раннем докембрии действовал механизм толстоплитовой тектоники, который к неопротерозою постепенно сменился на механизм тонкоплитовой тектоники. Мантийно-плюмовое воздействие на литосферу Земли - сквозное явление на протяжении всей геологической истории, которое определяет возникновение в ней существенных плотностных неоднородностей и, как следствие, мест инициации субдукции и роста континентальной коры.

Ключевые слова: бониниты; инициация субдукции; мантийный плюм; эволюция геодинамики Земли

The heavy is the root of the light. (Lao Tzu, Daodejing, 5th-4th century BC)

1. Introduction

Boninites attracted much attention at the end of the 20th century, primarily, due to deep-sea studies of forearc slopes in the modern plate convergence zones of the southwestern Pacific. At that time, however, detailed geochemical studies of many ophiolite sections revealed the presence of boninites unknown for mid-oceanic ridges, but abundant in juvenile island-arc basements. Following the common postulation that ophiolites are remnants of the ancient oceanic crust, ophiolites in fold belts were largely interpreted as relic sutures which mark sites of paleo-ocean closure. As the boninite findings were ever-growing in both the classical fragments of the 'ancient oceanic crust', e.g. the Troodos in Cyprus, and Semail in Oman, or 'classical sutures' such as, the Main Urals Fault in Russia, this inevitably raises the question on the ophiolite nature.

Another problem concerns the evolution of boninitic magmatism as boninites have been discovered in the early Precambrian greenstone belts. Until recently, it was believed that in the early stages of the Earth evolution, 'dry' komatiitic volcanism was predominant and then followed by 'wet' boninite volcanism [Hall, Hughes, 1993]. However, the komatiite abundance in the Archean seems to be well overestimated. Ko-matiites are lacking in many belts and constitute less than 5 % of the volcanic rocks in most Archean greenstone belts [de Wit, Ashwal, 1997]. Detailed geochemical studies of mafic-ultramafic complexes in greenstone belts show that a majority of olivine spinifex-free 'ko-matiites' and 'komatiitic basalts' belong in fact to the

boninite series rocks. According to [Furnes et al., 2014], boninitic magmatism has been occurred through the geological history of the Earth. However, aspects of its evolution have not been thoroughly discussed in the literature yet.

Thus, the aims of the paper are to provide an overview of the boninitic magmatism occurrences through space and time, track its evolution, and propose a geo-dynamic explanation of this highly informative phenomenon.

2. Definition and geochemical characteristics

of boninites and the boninite series

The International Union of Geological Sciences (IUGS) classification of the high-Mg volcanic rocks defined a boninite as a volcanic rock with the following arbitral chemical composition recalculated to 100 wt %: SiO2>52 wt %; MgO>8 wt %, and Ti02<0.5 wt % [Le Bas, 2000]. The Ti content is strongly constrained in the classification as a critical requirement because of titanium is the most incompatible among the major elements, and its low content in primitive melts indicates that their mantle source was highly depleted. Boninite definitions in many publications are very similar to the above-mentioned one. According to [Crawford et al., 1989], SiO2 content in boninite lavas exceeds 53 wt %, and Mg#>0.6, where Mg#=Mg/[Mg+Fe2+]. It was recommended in [Taylor et al., 1994], to classify volcanic and hypabyssal rocks as boninites if SiO2>53 wt %, Ti02<0.6 wt %, and 7<MgO<25 wt %.

The above definitions of boninites are based on chemistry, rather than on mineralogy of these rocks. Neither mineralogical nor petrographic characteristics could be included in the classical definition of boninites as the modal mineralogy of phenocrysts show highly variable compositions, and their textural patterns vary widely also. The traditional approach distinguishes between two major groups of boninites, high-Ca and low-Ca [Crawford et al., 1989], characterized by Ca0/Al203>0.75 and <0.6, respectively. Boninites with intermediate ratios are grouped as low-Ca type 3 boninites. The upper pillow lavas of the Troodos (Cyprus) ophiolite are distinguished as the tectonic and pet-rographic type of high-Ca boninites. Low-Ca type 1 boninites are best represented by the Cenozoic lavas of New Caledonia. Low-Ca type 2 boninites include high-Mg andesite lavas of the Baja peninsula, California and the Shikoku Island, Japan. The latter of very special geochemical characteristics and tectonic settings have their own names, bajaites and sanukites, respectively [Rogers, Saunders, 1989; Tatsumi, Ishizaka, 1981]. Low-Ca type 3 boninites are most fully represented in the Bonin Islands, Japan where the boninites have been described for the first time at the end of the 19th century.

Despite the fact that two clearly different geochemi-cal groups (high-Ca and low-Ca boninites) are recognized, geodynamic settings for particular types of bo-ninitic magmatism occurrences have not been specified yet [Meffre et al., 1996]. The assemblage including both high-Ca and low-Ca boninites and intermediate-Ca bo-ninites have been reported from the Mariana island arc [Arculus et al., 1992]. In other occurrences, such as the Troodos, Oman and Josephine ophiolites, only high-Ca boninites are present [Cameron, 1985; Crawford et al., 1989; Ishikawa et al., 2002; Harper, 2003]. Low-Ca and intermediate boninites are dominant in the Ordovician ophiolites of Newfoundland [Bedard, 1999] (Table 1).

Boninites are spatially and genetically related to primitive island-arc low-Ti lavas, thus generating a need for recognition of a separate magmatic series known as 'the boninite series' [Pearce, Robinson, 2010; and others]. The less differentiated rocks of the series are known from the literature as low-Ti tholeiites, LOTI, [Brown, Jenner, 1989] or low-Ti ophiolite basalts [Sun, Nesbitt, 1978], and they are typical of numerous ophiolites. In many cases, the boninite series volcanic rocks are represented mainly by primitive lavas (MgO>12 wt %) that can be termed as picrites or mistakenly classified as komatiitic basalts or basaltic ko-matiites [Cameron et al., 1979].

Crystallization of low-Ti tholeiite melts can be schematically given as follows: olivine ^ clinopyroxene ^ plagioclase, and the scheme for MORB melts is: olivine ^ plagioclase ^ clinopyroxene [Cameron et al., 1980; Natland, 1981]. This difference is due to drastic distinc-

tions between water-saturated melts of the boninite series and dry melts of the komatiitic and tholeiitic series. The komatiitic series trend is generated by dry mantle melting and directed towards a field of MOR tholeiites, while the boninite series trend reflects the evolution of the primary composition towards the quartz apex, which is typical of wet melting at gradually decreasing temperature and pressure. Distinctions in the compositional trends of the high-Mg volcanic series are well depicted in the projection of the normative basalt tetrahedron Ol-Pl-Qt plotted by using the data on the boninite series of the North Karelian greenstone belt and the komatiitic series of the Kostomuksha greenstone belt in the Baltic Shield (Fig. 1).

The trace elements patterns of the boninites demonstrate strong depletion of the mantle source and concurrent evidences of their suprasubduction genesis, such as, for instance, negative anomalies of Nb(Ta) (Fig. 2). Different types of the boninites have spectra of the same type with distinct negative anomalies of Nb(Ta) and Ti, positive anomalies of Sr and Zr(Hf) and apparent enrichment in large-ionic lithophile (LIL) elements (Rb, Ba, Cs, U, and Th) relative to N-MORB. Such a kind of the trace elements patterns clearly suggests that (i) petrogenesis of boninite series should be excluded any and even a minimum crustal input, and (ii) melting requires a mantle source more depleted in comparison to the mantle lherzolite that generates MORB melts. In other words, melts of the boninite series seem to be derived from a depleted harzburgite mantle source [Hickey, Frey, 1982; Crawford et al., 1989].

On the other hand, there is geochemical evidence that there is a direct link between the boninite series rocks and island-arc tholeiites (IAT), taking into account the commonly accepted idea that juvenile continental crust portions are generated in the island-arc zones (Fig. 3). Thus, the boninite series rocks are of crucial importance for deciphering crust-forming processes through the geological history of the Earth.

3. Boninites in the geological time and space

The global occurrences of boninite series rocks are shown on Figure 4 and at Table 2. Boninites have been discovered in all the continents, except South America, which is attributable to the fact that South America is still relatively poorly covered by geological studies. The boninite series rocks are described in various sequences through the geologically documented history of the Earth, from the most ancient Eoarchean Isua belt in the southwestern Greenland to the recent eruptions of submarine volcanoes in the northeastern Lau basin. The only lacuna in their history is the Mesoproterozoic, the period of relatively passive tectonic events, possibly

Table 1. Selected analyses representative of boninite series of different ages which are mentioned in the text and were used for the geochemical plots

Таблица 1. Представительные анализы пород бонинитовой серии разного возраста, упоминаемые в тексте и использованные для построения диаграмм

Location Northern Tonga Izu-Bonin-Mariana Forearc Troodos Ophiolite Dunzhugur Ophiolite

Age 2.5-2.7 Ma 48-42 Ma 90-95 Ma 1020 Ma

Source Falloon et al., 2008 Pearce et al., 1992 Cameron, 1985 Sklyarov et al, 2016

Type HCB HCB LCB ICB HCB HCB LOTI LOTI LCB ICB LOTI

Si02 (wt.%) 53.32 55.59 53.47 53.00 47.23 53.80 49.81 51.74 58.09 56.69 48.75

Ti02 0.24 0.22 0.20 0.25 0.31 0.31 0.19 0.19 0.18 0.18 0.14

AI2O3 11.45 14.87 12.05 12.46 12.06 11.40 11.22 12.47 14.41 9.97 16.94

FeO* 9.10 8.25 7.62 6.78 7.24 7.52 8.08 7.20 6.96 7.91 7.45

MnO 0.16 0.13 0.13 0.14 0.14 0.16 0.15 0.15 0.15 0.17 0.14

MgO 13.37 9.19 13.04 13.08 11.40 12.54 13.92 12.74 9.46 13.5 8.9

CaO 8.85 8.92 5.27 6.70 11.90 8.73 10.12 9.88 5.63 6.34 11.04

NazO 2.66 1.83 3.19 2.76 2.12 1.30 0.66 1.10 1.46 0.37 2.14

K2O 0.93 0.80 0.67 0.21 0.26 0.15 0.21 0.39 2.78 3.98 0.78

P2O5 0.02 0.03 0.03 0.04 0.04 0.02 0.02 0.01 0.08 0.11 0.02

Mg# 73.4 67.6 76.2 78.4 74.7 76.9 78.4 77.8 71.8 76.2 6.92

Cr (ppm) 812 477 1036 786 762 910 980 940

Ni 228 124 306 296 374 268 378 339

Co 45 38.5 44.6

Sc 44 48 29 33.7 27.1 37 40 40

V 265 295 166 202 193 212 212 203

Ba 119 112 47 7.6 15 12 18 45 273 504 127

Rb 5.7 20.3 6.4 3.1 7.6 3 7 7 61.2 68.2 6.8

Sr 118 74 107 136 139 62 53 62

Nb 1.04 0.96 0.53 0.44 0.53 1 0.8 1 1.16 1.98 0.69

Hf 0.39 0.39 0.9 0.70 0.70 0.7 0.2 0.3 0.82 0.77 0.3

Zr 12.0 12.1 35 27 26 18 10 16 19.9 26.9 10

Y 4.7 4.9 7.2 6.5 8.2 7.3 7 7 7.71 7.7 4.56

Th 0.37 0.37 0.27 0.17 0.18 0.67 0.9 0.3

U 0.19 0.14 0.18 0.12 0.14 0.17 0.44 0.25

La 2.14 1.94 1.81 1.15 1.50 0.88 0.72 1.09 3.59 3.31 1.26

Ce 4.12 3.97 3.86 2.57 3.05 2.22 1.46 2.19 7.77 7.11 2.72

Pr 0.53 0.52 0.59 0.37 0.50 0.36 0.15 0.23 0.76 0.89 0.33

Nd 2.32 2.32 2.84 1.97 2.61 1.91 0.58 1.02 3.41 3.83 1.31

Sm 0.62 0.62 0.81 0.61 0.80 0.73 0.23 0.32 0.94 0.82 0.36

Eu 0.22 0.26 0.27 0.22 0.31 0.27 0.10 0.11 0.24 0.3 0.96

Gd 0.22 0.76 0.97 0.81 1.06 0.96 0.44 0.51 0.94 0.87 0.52

Tb 0.12 0.12 0.16 0.15 0.20 0.20 0.10 0.11 0.17 0.14 0.12

Dy 0.84 0.89 1.05 1.00 1.30 1.39 0.90 0.93 1.06 1.08 0.72

Ho 0.18 0.19 0.23 0.31 0.30 0.32 0.23 0.23 0.22 0.23 0.17

Er 0.56 0.61 0.69 0.72 0.84 1.03 0.76 0.76 0.84 0.66 0.51

Tm 0.09 0.09 0.11 0.12 0.14 0.12 0.12 0.08

Yb 0.61 0.62 0.80 0.75 0.87 1.04 0.84 0.84 0.78 0.77 0.58

Lu 0.10 0.11 0.14 0.11 0.13 0.23 0.13 0.09

Table 1 (end) Таблица 1 (окончание)

Location Flin Flon belt Abitibi greenstone belt North Karelian greenstone belt Isua greenstone belt

Age 1900 Ma 2715 Ma 2800 Ma 3700- 3800 Ma

Source Wyman, 1999b Kerrich et al., 1998 Shchipansky et al, 2004 Polat et al., 2002

Type LCB HCB ICB LOTI LOTI LOTI ICB НСВ LOTI LCB

SiOz (wt%) 53.8 57.6 59.7 43.8 50.5 49.16 52.80 57.52 49.08 54.00

Ti02 0.12 0.09 0.31 0.18 0.24 0.42 0.47 0.35 0.23 0.40

AI2O3 14.50 7.6 15.8 15.8 14.1 10.65 13.03 9.39 16.94 17.74

FeO* 9.81 9.27 6.90 10.1 9.70 11.03 10.90 9.53 9.12 8.37

MnO 0.10 0.2 0.50 0.21 0.27 0.19 0.17 0.17 0.17 0.13

MgO 11.1 12.6 7.37 22.0 15.0 12.92 10.35 12.80 15.57 7.05

CaO 5.0 9.6 6.59 6.7 6.18 11.37 7.37 6.86 6.82 8.39

NazO 3.7 0.6 1.98 0.12 2.48 0.35 2.44 0.07 0.95 2.85

K2O 0.8 1.5 0.01 0.01 0.01 0.07 0.08 0.10 0.09 0.12

P2O5 0.01 0.02 0.06 0.03 0.03 0.13 0.10 0.08 0.01 0.02

Mg# 68.0 71.9 69.0 80.0 74.0 0.71 66.0 71.8 77.0 63.0

Cr (ppm) 544 1411 330 1460 1058 706 324 540 951 141

Ni 155 183 210 760 499 254 187 189 376 69

Co 65 59 51 74 71.3 74 53 51 63 35

Sc 85 49 55 48 44 38 40 33 42 45

V 319 170 250 250 160 238 339 202 192 195

Ba 15.8 18 12 8 14

Rb 0.92 1.2 2.7 7.1 4.3

Sr 49 63 69.5 54 116

Nb 1.2 0.6 1.3 0.42 0.71 0.88 1.01 1.06 0.15 0.55

Hf 0.71 0.35 0.64 0.3 0.38 0.89 0.83 0.75 1.1

Zr 26 14 26 12 16 34 29 26.5 16.4 27.5

Y 5.0 3.0 17 9 10 12.2 13.9 0.17 9.2 13

Th 0.36 0.2 0.2 0.05 0.09 0.09 0.13 0.08 0.08 0.23

U 0.31 0.16 0.01 0.04 0.04 0.02

La 1.78 1.21 1.46 0.49 0.72 0.78 1.16 1.39 0.46 1.26

Ce 4.11 2.67 3.87 1.33 1.81 2.51 3.49 3.55 1.26 3.08

Pr 0.51 0.31 0.59 0.21 0.27 0.39 0.55 0.52 0.20 0.46

Nd 2.18 1.32 2.62 1.11 1.33 2.05 3.42 2.57 1.10 2.28

Sm 0.58 0.33 0.74 0.39 0.45 0.73 0.99 0.82 0.44 0.91

Eu 0.23 0.12 0.19 0.14 0.24 0.35 0.39 0.29 0.13 0.28

Gd 0.76 0.45 1.19 0.63 0.67 1.31 1.62 1.37 0.69 1.32

Tb 0.13 0.07 0.26 0.14 0.14 0.24 0.27 0.27 0.16 0.26

Dy 0.89 0.43 2.54 1.32 1.22 1.77 2.03 1.78 1.33 2.04

Ho 0.19 0.1 0.65 0.36 0.31 0.41 0.45 0.43 0.37 0.54

Er 0.55 0.29 2.33 1.23 0.95 1.20 1.32 1.26 1.36 1.91

Tm 0.10 0.05 0.39 0.22 0.17 0.17 0.18 0.19 0.21 0.28

Yb 0.55 0.3 3.03 1.49 1.19 1.05 1.24 1.18 1.49 1.83

Lu 0.10 0.05 0.50 0.24 0.18 0.15 0.19 0.19 0.24 0.28

Note. НСВ - High-Ca boninite, ICB - intermediate-Ca boninite, LCB - Low-Ca boninite; LOTI - Low-Ti basalt/picrite. Subdivision of boninites is after [Crawford etal., 1989].

Примечание. НСВ - высококальциевые бониниты, ICB - умереннокальциевые бониниты, LCB - низкокальциевые бониниты, LOTI - низкотитанистые базальты/пикриты. Разделение бонинитов по [Crawford etal., 1989].

PI

MORB glasses

From Di

: * v- • •

■.v • v

p

+ Komatiite series • Boninite seriesN

01

Opx

Qz

Fig. 1. A comparison between normative compositional trends of the boninite and komatiite volcanic series in ternary projection olivine-plagioclase-quartz from diopside [Walker et al., 1979]. The both volcanic series are from the Archean Karelian Craton. Modified after [Shchipansky, 2008].

Рис. 1. Положение фигуративных точек состава позд-неархейских высокомагнезиальных вулканитов Карельской ГЗО в проекции нормативного базальтового тетраэдра Ol-Pl-Qz из вершины диопсида (Di) [Walker et al., 1979]. Обе вулканические серии из архейского Карельского кратона. Модифицировано из работы [Shchipansky, 2008].

related to the existence of the stable supercontinent Nuna [Cawood, Hawkesworth, 2014].

Contemporary settings for generation of the boninite series rocks are known in the areas of intraoce-anic island-arc systems of the southwestern Pacific, specifically the Izu-Bonini-Mariana (IBM) and Tonga-Kermadec island arcs. Boninites compose mostly fore-arc slopes, forming the lowest stratigraphic levels in the island-arc architecture. These regions are natural laboratories for comprehensive studies of the boninitic magmatism coupled with geodynamic modeling of its genesis. It is noteworthy that since the Paleocene onwards there is a strong link between boninites and ophiolites, as confirmed by the data on the Cape Vogel Peninsula, Papua New Guinea (see Table 2). Actually, this genetic link has long been known [Sun, Nesbitt, 1978; Cameron et al., 1979] and led to introduction of the term 'suprasubduction zone (SSZ) ophiolite' in the 1980s to acknowledge that some ophiolites are more closely related to island arcs than to ocean ridges [Pearce, 1982]. This term was readily accepted in the literature, and the majority of the known ophiolite complexes, including the largest ones with the com-

pletely preserved sequences, e.g. Troodos, Oman, were classified into the SSZ type.

According to [Sklyarov et al., 2016], there are four main types of boninites in the ophiolite sequences (Fig. 5). Type 1 bonninites spatially coexists with ophi-olites, although compose (or belong to) other tectonic units. Type 2 boninites presents as later constituents of ophiolite sequences, such as crosscutting dikes or lavas on top of section. Type 3 includes island-arc tholeiites and basaltic andesites coupled with boninites, which are replaced by younger MORB or BABB affinities. Type 4 boninites occupies the whole mafic portion of ophio-lite sequences, together with island-arc tholeiites and basaltic andesites, and all the components of such sequences (gabbro, dykes, and lavas) have clearly bo-ninitic affinities.

These types of the boninite-bearing ophiolite sequences are recorded through the entire Phanerozoic and Neoproterozoic. The Neoproterozoic boninite

50

ш

10

Ш >

a.

Q.

о о a.

0.5

1 1 1 1 1 1 1 1 1 1 1 1 1 II 1 1 1 1 N-MORB

- I j\ high-Ca boninite

low-Ca boninite

- \i у - " -

~ 1 1 1 1 1 1 1 1 1 1 1 medium-Ca boninite ~ 1 1 1 1 1 1 1 "

Rb Ba U Th N b LaCeSr PrNd ZrHfSm EuGdTi TbDyHoY ErTm Yb

Fig. 2. Trace elements patterns of different boninite groups from Izu-Bonin arc by comparison with N-MORB.

Note the negative anomalies of Nb (Ta) and, conversely, positive anomalies of Zr (Hf) and Sr, both typical of supra-subduction zone volcanics. Strong depletion in REE and HFSE of the boninite compositions with respect to N-MORB is clearly visible. Averaged boninite compositions are from [Pearce et al., 1992]. N-MORB and primitive mantle values are from [Hofmann, 1988].

Рис. 2. Мультиэлементные диаграммы разных геохимических групп бонинитов Изу-Бонинской островной дуги.

Заметим характерные для бонинитов и индикаторные для петрогенезиса вулканитов в надсубдукционных обстановках отрицательные аномалии Nb (Ta) и Ti и, наоборот, положительные Zr (Hf) и Sr. Ясно видны также геохимические различия в уровне общей деплетации некогерентными элементами между бонинитами и толеитами MORB. Составы бонинитов по [Pearce et al., 1992]. Примитивная мантия и MORB по [Hofmann, 1988].

200

100 -

Enriched LILEs

Enriched K, Pb, and Sr

Depleted Nb, Та

\ Boninite __

"Depleted HREE

~r

~l-г

"Г

"Г

CsRbBaTh U Nb Та К La CePb NdSrSmHf Zr Ti EuGd Dy Y ErYbLu Na

Increasing compatibility

Fig. 3. Spidergram for trace-element compositions of the continental crust and juvenile arc melts (boninite, arc tholeiite), modified after [Stern, 2002].

Elements on the horizontal axis are listed in order of their incompatibility in the mantle relative to melt; elements on the left are strongly partitioned into the melt whereas those on the right are strongly partitioned into a peridotite source. Note the characteristic enrichments of subduction zone outputs relative to MORB with respect to fluid-mobile LIL elements and the relative depletion of these in HFSE, and HREE. Note also the overall similarity of continental crust to juvenile arc crust.

Рис. 3. Спайдерограмма распределения малых элементов в континентальной коре и в расплавах ювенильных островных дуг (бонинита и островодужного толеита), модифицировано из работы [Stern, 2002].

По горизонтальной оси графика элементы расположены в порядке их несовместимости в мантии относительно расплавной фазы. В левой части спектров расположены элементы, совместимые с расплавами; в правой части - совместимые с перидотитовым источником. Заметим характерное для субдукционных зон обогащение мобильными крупноионными элементами, связанное с флюидным привносом, и, наоборот, деплетацию высокозарядных ионов и тяжелых редких земель по сравнению с N-MORB. Обращает на себя внимание общее сходство геохимии континентальной коры и ювенильной коры островных дуг.

series are widely distributed in the Altai-Sayan area of the southern frame of the Siberian Craton [Simonov et al., 1994; Dobretsov et al., 2005]. More ancient boninite series are discovered in the Archean and Paleoprotero-zoic metamorphosed fold belts that are often jointly termed as 'greenstone belts', regardless of the meta-morphic grade.

As a rule, the rock assemblages composing the greenstone belts were strongly tectonically transformed and dismembered that gives a zero chance for preservation of a complete ophiolite sequence. Nonetheless, the Early Precambrian greenstone belts maintain vestiges indicating that rock assemblages of oceanic provenance were involved in the tectogenesis of the belts [Shchipansky, 2008; Rosen et al., 2008; Furnes et al., 2015]. These are mainly the isotopic and geo-chemical signatures of the lack of crustal contamination or affiliation of the mafic-ultramafic assemblages of the greenstone belts to the ensimatic island-arc settings. In the latter case, the presence of the boninite series rocks is a critical support to the suprasubduction ophiolite interpretation of the greenstone belts sequences.

Besides the isotope-geochemical data, some boni-nite series from the Archean greenstone belts do preserve field evidences on the ocean lithosphere extension. We reported a fragment of the suprasubduction ophiolites with sheeted dikes in gabbroids and me-talavas of the boninite series (~2.8 billion years) in the Iringora locality of the North-Karelian greenstone belt [Shchipansky et al., 2001, 2004]. Later on, relics of a sheeted-dike complex were discovered in the Garbenchifer formation from the Eoarchean Isua su-pracrustal belt, southwestern Greenland [Furnes et al., 2007, 2009]. This suggests that the boninite series rocks and the SSZ spreading have been genetically related since the early stages of the Earth's geological history.

The early Precambrian boninite series differ from the Phanerozoic ones by the presence of both boninites and komatiites in some greenstone belts, such as the Bogoin (Paleoproterozoic), Abitibi (Neoarchean), and Koolyanobbing (Mesoarchean) belts, which is related to mantle plume fingerprinting into intra-oceanic convergence zones (see Table 2). It should be noted, however, that data on the Phanerozoic boninite series often

North Karelian (2.8 Ga)

Ballantrae ¿ZP (500-480)

S. Anyui—о (280-270).,

Isua (3.7-3.8 Ga)

Thetford Mines - Bett Coves (500-480)

Yukon-Tanana (360-350)

Magnitigorsk

,'Charsk Dunzhugur (1.02 Ga) (380-320) О Altai-Sayan (600К~Л

—Ч^о^-^о^ С У ^Mongolia\V\(

—N North Qilin^T^ \ (505-490) \ Jiangnan Л \(0.8 Gay 1

Frotet-Evans (2.8 Ga) Abitibi (2.7 Ga)

Josephine (164-162)

Mirdita'

Bangong (200-150;

Coast Range (172-166)

Izu-Bonin (48-43)

Troodos' (95-90)

Oman , .(96-93)

Mariana (48-43)

Gadwal (2.7 Gab js¡nghbhum] ^ V (3.5-3.4 Ga)

Adola (0.8-0.9 Ga) ^

Zambales (48-44)/)

Bogoin (2.3 Ga)l

Cape Vogel (65-55)

Lau

\(1-0)

Koolyanobbing (3.0 Ga)

Heathcote (515-510V

Meso-Cenozoic

Western Tasmania (515-510)

Phanerozoic

Late Precambrian

Victoria Land

Early Precambrian

| Fig. 4. Global map of boninite occurrences from which boninite series rocks have been reported and described in details. See Table 1 for references.

о

(С

о а.

с

0

01

3

о (л

й»

о

3 О ■и

<

о с 3

1С -ч

(Я (Л

с

(Б ю

Рис. 4. Распространение бонинитов на земном шаре, детальное описание находок которых было приведено в литературе. См. таблицу 1, где приведены соответствующие ссылки. л

(Л

Table 2. Global occurrences of boninite series volcanics with some short comments on their tectonic characteristics

Таблица 2. Распространенность вулканитов боннннтовой серии на земном шаре и краткая характеристика областей их локализации

Locality

Age

Tectonic setting

Geodynamic implication

References

NE Lau Basin

Northern termination of the Tonga trench

Yap trench, Philippine sea

Izu-Bonin-Mariana Island Arc

Luzon, Philippines

Cape Vogel Peninsula, PNG

Troodos Massif, Cyprus

Pindos Ophiolite Complex, NW Greece

Mirdita Ophiolite Complex, northern Albania

Josephine Ohiolite, Klamath Mountains, NW California

Coast Range ophiolites, California

Modern submarine explosions

2.5-2.7 Ma

7-8 Ma 48-42 Ma

48-44 Ma Paleocene 90-95 Ma

170-165 Ma Jurassic 164-162 Ma

172-166 Ma

Bangong Lake ophiolite, NW Early Jurassic Tibet

South Anyui suture ophiolite, 280-270 Ma Western Chukotka

Multi-spreading centre Fore-arc slope

Trench slope Fore-arc slope

Complete SSZ ophiolite sequence Papuan Ultramafic Belt Complete SSZ ophiolite sequence

Complete SSZ ophiolite sequence Complete SSZ ophiolite sequence Complete SSZ ophiolite sequence

Ophiolite remnants in serpentinite matrix mélange

Ophiolite remnants in serpentinite matrix mélange

Gabbro, gabbro-pyroxenite, and restitic ultra-mfic rocks of an ophiolite sequence

Nascent juvenile island arc above a subduction zone

Supra-subduction extensional zone in an island arc basement

Supra-subduction extensional zone in a forearc setting

Nascent juvenile island arc above a subduction zone

Basement of an juvenal island arc

Supra-subduction extensional zone in a forearc setting

Nascent juvenile island arc above a zone of subduction initiation

Supra-subduction extensional zone in a forearc setting

Supra-subduction extensional zone in a forearc setting possibly related to the MOR subduction

Progradation of a spreading centre from backarc basin to the forearc area of the juvenile island arc

Supra-subduction extensional zone in a forearc setting possibly related to the MOR centers subduction

Supra-subduction extensional zone in a forearc setting

Supra-subduction extensional zone in a forearc or backarc settings

Resing et al., 2011

Falloon et al., 2008; Sobolev, Danyushevsky, 1994; Falloon, Crawford, 1991

Crawford et al, 1986

Crawford et al., 1981; Crawford et al, 1989; Stern et al, 1991; Stern, Bloomer, 1992; Stern, 2004; Taylor et al, 1994; Li et al, 2013

Yumul, 1996; Yumul et al., 2000

Walker, Cameron, 1983; Crawford et al, 1989

Sun, Nesbitt, 1978; Cameron et al, 1980; Cameron, 1985; Sobolev et al, 1993; Pearce, Robinson, 2010; Whattam, Stern, 2011

Pe-Piper et al, 2004

Bortolloti et al, 1996; Bortolloti et al, 2002

Harper, 1984; Harper, 2003; Wallin, Metcalf 1998; Metcalf Shervais, 2008

Shervais, 2001; Shervais et al, 2004 Shervais et al, 2005

Shi et al., 2004; Shi et al, 2008

Ganelin, Silantjev, 2008

Table 2 (continued)

Таблица 2 (продолжение)

Locality Age Tectonic setting Geodynamic implication References

Koh Ophiolite, New Caledonia 295 Ma Upper lavas of the complete SSZ ophiolite Progradation of a spreading centre from backarc basin to the forearc area of the juvenile island arc Meffreet al., 1996; Aitchisonet al., 1998

Charsk belt, Eastern Kazakhstan Late Devonian -Early Carboniferous Mafic to acid volcanics associated with dismembered ophiolites Supra-subduction environment Simonov et al, 2010

Yukon-Tanana terrane, SW Yukon, Canada Pre Late Devonian Fire Lake boninitic rocks containing volcanic-hosted massive sulfide deposits Progradation of spreading ridge into a suprasubduction environment Piercey et al, 2001

Magnitogorsk juvenile island arc of the Sakmar zone in the South Ural Early Devonian Volcanics and dykes from the SSZ ophiolite sequences Supra-subduction extensional zone in a basement of juvenal arc Spadea et al., 1998; Kosarev et al, 2005; Puchkov, 2010; Belova et al, 2010

Ballantrae complex ophiolite, Southwest Scotland 500-480 Ma Volcanics from the SSZ ophiolite sequences Fragment of an intra-oceanic island arc Smellie et al., 1995

Thetford Mines Ohiolite, Canadian Appalachians 500-480 Ma Volcanics from the SSZ ophiolite sequences Supra-subduction extensional zone in a forearc setting Page et al., 2009

Betts Cove ophiolite, Newfoundland, Canada 500-480 Ma Volcanics from the SSZ ophiolite sequences Supra-subduction extensional zone in a forearc setting Bedardet al., 1998; Bedard, 1999

Ophiolite of Western Tasmania Early Cambrian Lavas associated with ophiolitic ultramafic complexes Supra-subduction extensional zone in a forearc setting Brown, Jenner, 1989; Crawford, Berry, 1992

Heathcote Greenstone belt in the Lachlan Fold Belt, SE Australia Early Cambrian The oldest metavolcanics in the Lachlan Foldbelt Supra-subduction extensional zone in an island arc setting Crawford, Cameron, 1985

Tholeiite-boninitic terraine of the North Qilian suture zone, China Early Cambrian The volcanic complex of the North Qilian ophiolite belt Nascent juvenile island arc above a subduction zone. Xia et al., 2012

Boninite-derived amphibolites from the Lanterman-Mariner suture, Victoria Land Cambrian - Ordovic Boninite-derived retrogressed amphibolites within the Early Paleozoic sure zone Juvenile island arc or backarc basin Palmeri et al, 2012

Boninite series of the Altai-Sayan Fold Belt Vendian-Cambrian Lavas, dykes and gabbro of boninitic affinity from dismembered or partially preserved SSZ ophiolite sequences Subduction initiation of the nascent juvenile island arcs Simonov et al, 1994; Dobretsov et al, 2005

Boninite-series rocks from the ophiolites of Western Mongolia Vendian-Cambrian (-600 Ma) Lavas, dykes and gabbro of boninitic affinity from the Khantaishir and Bayannor Ranges Ophiolites Forearc or backarc extensional basins Zonenshain, Kuzmin, 1978; Kepezhiskas et al, 1991; Khain et al, 2003

Boninite-series rocks from the Jiangnan Fold Belt, SW Cina 850-825 Ma Greenschist-facies boninite series rocks of the SSZ opniolite conplex from the Xiuning-Shexian Fault zone Nascent juvenile island arc above a subduction zone Zhao, Asimov, 2014

Table 2 (end)

Таблица 2 (окончание)

Locality Age Tectonic setting Geodynamic implication References

Adola boninite from the Megado ophiolitic belt, Southern Ethiopia 790 Ma Low-grade boninite series rocks of the Megado SSZ ophiolite Supra-subduction extensional zone in a forearc or backarc settings Wolde et al., 1996; Yibas et al., 2003

Dunzhugur boninite series of the Ilchir Ohiolite Belt, Eastern Sayan, Russia 1020 Ma Lavas, dykes and gabbro of boninitic affinity from the upper part of ophiolite sequence Nascent juvenile island arc above a subduction zone Dobretsov et al, 1986; Khain et al, 2002; Kuzmichev, 2004

Boninites of the Flin Flon greenstone belt, Trans-Hudzon Orogen, Canada 1900 Ma Low grade subaqueous lavas in the lowermost part of the island arc volcanics Supra-subduction extensional zone in a forearc or backarc settings Stern et al., 1995; Leybourne et al, 1997; Wyman, 1999a

Bogoin boninite-like rocks of the Bogoin-Boali greenstone belt, Central African Republic 2300 Ma Upper pillowed amphibolites underlined by Al-depleted and Al-undepleted komatiites and meta-greywackes Accretionaiy complex in a convergent plate margin Poidevin, 1994

Boninites from the Gadwal grrenstone belt, Eastern Dharwar Craton, India 2700-2500 Ma Amphibolite-facies boninite series occurred along with pillowed metatholeiites Subduction environments of small plate tectonics involving an interaction of mantle plume and inta-oceanic island arc Manikyamba et al, 2005; Manikyamba, Kerrich, 2012

Boninite series of the Abitibi greenstone belt 2715 Ma Greenschist-facies volcanics Wihtney underlined by Al-depleted and Al-undepleted komatiites Kidd-Munro Accretionaiy complex emerged by interaction of mantle plume and inta-oceanic island arc Kerrich et al., 1988; Wyman, 1999a; Wyman, 1999b

Boninite series volcanics of the Frotet-Evans greenstone belt, Opatica, Superior, Canada 2790 Ma Low-grade boninitic ricks of the lowermost part of the arc-related Assinica group sequence Supra-subduction extensional zone in a forearc setting Boily, Dion, 2002

Boninite series of the North Karelian greenstone belt, Baltic Shield, Russia 2804 Ma Amphibolite-facies boninite-series rocks from the upper part of the SSZ ophiolite complex Nascent juvenile island arc above a subduction zone Shchipansky et al, 1999; Shchipansky et al, 2001; Bibikova et al, 2003; Shchipansky et al, 2004 Shchipansky, 2008

Boninite series of the Koolyanobbing greenstone belt, Craton Yilgarn, SW Australia 3023 Ma Amphibolite-facies boninite-series rocks alternated with komattite-tholeiite series volcanics in the lower part of the Koolyanobbing sequence Accretionaiy complex emerged by interaction of mantle plume and intra-oceanic island arc Angerer et al, 2013

Boninite series of the Older Metamorphic Group of the Singhbhum Craton, eastern India 3550-3440 Ma Amphibolite-facies boninite-series rocks associated with metapelites and BIFs One of the first evidences on a subduction initiation at the Archean Manikyamba et al, 2015

Boninite-series of the Isua greenstone belt, SW Greenland) 3800 Ma Amphibolite-facies boninite-series rocks of Garbenchifer Formation associated with IAT and BIFs. First evidence on a subduction initiation and oceanic lithosphere extension in the Earth histoiy Polat et al., 2002; Polat, Hoffman, 2003; Fumes et al, 2007; Fumes et al, 2009

1 type

2 type

3 type

4 type

MORB+BABB

IAT

IAT+boninite

boninite

harzburgite, dunite

pyroxenite layered gabbro massive gabbro sheeted dykes cutting dykes

П П П

г г г г

lavas, mostly pillowed

Г Г Г

п п п

г г г г

± ±

_L ± ± 1

п п п

г г г г

Fig. 5. Models illustrating four tectonic types of boninite occurrences in ophiolitic sections, after [Sklyarov et al., 2016]. Рис. 5. Идеализированные разрезы «тектонических» типов бонинитов, по [Sklyarov et al., 2016].

suggest the mantle plume involvement in their petro-genesis that will be discussed below.

4. Petrogenesis of the boninite series rocks, and their evolution in time

The boninite series rocks are unique because of their genesis requires the set of conditions which is possible only in specific geodynamic settings within spatially limited locations. Indeed, the origin of a bo-ninite source requires prior depletion of an upper mantle reservoir by basalt melt extraction in one or several episodes, which means that mantle harzburgite was such a source [Hickey, Frey, 1982; Sun, Nesbitt, 1978; Duncan, Green, 1980, 1987]. Boninitic melts are primitive and, at the same time, highly silicic with very low absolute concentrations of incompatible trace elements (Nb, Ta, and Ti) and rare earth elements; this suggests melting of residue peridotite at relatively shallow depths.

In fact, contents of Cr and Ni in the boninite series rocks are high. The strong and fast depletion in Cr suggests that chromite was among the main liquidus phases at the early stages of the primary melt fractionation. Depletion in Ni and simultaneously increasing contents of SiO2 indicate that olivine played an important role in magmas fractionation to Mg#~0.70. As noted above, the subsequent crystallization phase was clinopyroxene replaced then by plagioclase.

The boninite series lavas are considerably enriched in LILE and LREE in comparison to HFSE. Such geo-chemical characteristics suggest active infiltration of hydrous fluid into the mantle source of the boninite melts [Pearce et al., 1984, 1992; Pearce, 1982, 2003; Crawford et al., 1989; Bloomer et al., 1995; Stern et al., 1991]. The presence of fluid water in the mantle wedge is also a prerequisite to lower the solidus melting temperature of a refractory source.

While the ideas concerning the source of the bo-ninite series melts are commonly consistent, there is a disagreement in estimated P-T conditions for genera-

tion of parental magmas, i.e. primitive high-Mg and low-silica melts [Cameron, 1985]. In [Duncan, Green, 1980, 1987], the MgO content is estimated at 15-16 % for the primary Troodos boninite series melts generated by mantle harzburgite melting at 7-8 kbar at a temperature of 1360 °C. In [Sobolev et al., 1993], P-T estimates show larger depths and higher temperatures: 2-3 GPa; 1380-1450 °C; MgO content of 17-22 wt %. Resulting from the melt inclusion studies, the Neopro-terozoic boninite series from the Altai-Sayan segment of the Central-Asian fold belt were crystallized at T~1330-1150 °C from a water-saturated (up to 3.9 wt % H2O) primary melt at depths of about from 60 to 90 km [Dobretsov et al., 2005].

It has been experimentally determined [Kushiro, 1990] that adding 4.4-6.6 wt % H2O into the mantle substance reduces the melting start pressure by almost 2 kbar and lowers the temperature by ~150 °C. Experiments with refractory harzburgite under anhydrous and H2O-undersaturated conditions [Falloon, Danyu-shevsky, 2000] show that the high-Ca boninite pedogenesis requires temperatures as high as ~1480 °C and a pressure of about 1.5-2 GPa in the presence of 2-3 wt % H2O.

Notwithstanding significant uncertainties in the boninite series melt models (which compositions are dependent on many factors, including the mantle depletion degree and the fluid melting mode), there is a general consensus that their genesis requires anomalously high temperatures and the presence of a water-containing fluid in considerable quantities. In this case, temperature anomalies mean temperature values considerably exceeding those in the ambient upper mantle, which are sufficient for MORB generation. In the recently introduced theory on decompression melting of the upper mantle, a potential mantle temperature (Tp), i.e. a temperature at the lowest point of the mantle column, is set at ~1350 °C for generation of the parental melt for N-MORB [Langmuir et al., 1992; O'Hara, Herzberg, 2002]. This melt corresponds to a picrite composition (13 wt % MgO), and MORB13 composition model is perfectly matched to the fractionation trend of tholeiitic melts forming N-MORB [Mu, O'Hara, 2009]. Figure 6 illustrates the idea that the generation conditions of the boninite series rocks and MORB are different. Comparison of the MORB and boninite compositions with similar Mg -number gives grounds to challenge the geodynamic model suggesting that boninites could form in zones of stretching above the mid-ocean ridges in supra-subduction environments (see also [Metcalf, Shervais, 2008]). What petrogenetic conditions may be responsible for the boninite series rock genesis? Did any evolutionary changes occur in such conditions through the Earth's history?

Answers to the above questions can be derived from the theory of upper mantle decompression melting and

Fig. 6. A comparison between trace element compositions of the parental melt for modern MORB expressed as MORB13 [Niu, O'Hara, 2009] and the recent boninite with ~13 wt. % MgO from the Tonga fore-arc [Falloon et al., 2008].

It is clear that a petrogenesis of boninite series seems to be unrelated with an extensional environment characteristic of a MORB-forming magmatic column. Primitive mantle values are from [Hofmann, 1988].

Рис. 6. Сравнение состава малых элементов родона-чального расплава современных MORB, выраженного как MORB13 [Niu, O'Hara, 2009], и современного бони-нита с содержанием MgO ~13 вес. % из преддуговой области Тонга [Falloon et al., 2008].

Из рисунка следует, что петрогенезис бонинитовой серии не может быть связан с обстановками растяжения, в которых образуются расплавные колонны толеитов MORB-типа. Примитивная мантия по [Hofmann, 1988].

its key concept of accumulated fractional melting. If the latter is the case, the mantle source begins to melt in small amounts during decompression; typically, melt is extracted after 1-2 % melting by buoyancy-driven porous flow; and the residue continues to melt in small increments during decompression. This process takes place repeatedly as decompression progresses from an initial high to a final low melting pressure. The small melt fractions mix in transitional crustal magma chambers to make an aggregate primary magma melt. Each melt droplet contributing to the aggregate is in equilibrium with a specific source which composition varies from the initial relatively fertile to the depleted final composition. An aggregate fractional melt is not in equilibrium with its residue; only the final drop of the melt extracted is in equilibrium with the residue [Herzberg, Rudnick, 2012]. In solutions of the mass balance equation, mean values of residue, pressure and melt amount are used for simplicity. This facilitates solving

Fig. 7. A plot illustrating possible relationships between mantle potential temperatures (Tp) and primitive magma liquidus temperatures (Tliq) for the modern boninite series (Tonga, Troodos) and the Archean ones (Abitibi, North Karelian greenstone belt).

The diagram is simplified by showing a dry solidus applicable to both MORB and the boninite series, because a hydrous melting will take place below the dry solidus and the boninite series have a more depleted mantle solidus than MORBs. Tp and Tliq for boninite series rocks have been calculated from the appropriated compositions using software PRIMELT2.XLS [Herzberg, Asimow, 2008]. Data source: Troodos [Pearce, Robinson, 2010]; Tonga [Falloon et al., 2008]; Abitibi [Kerrich et al., 1998]; North Karelian greenstone belt [Shchipansky et al., 2004]. Data for the Kostomuksha komatiites were taken from [Puchtel et al., 1998].

Рис. 7. Иллюстрация соотношений между мантийными потенциальными температурами (Тр) и ликвидусными температурами (Tliq) примитивных магм современных (Тонга, Троодос) и архейских (Абитиби, Северо-Карельский зеленокаменный пояс) бонинитовых серий.

Диаграмма показывает упрощенные соотношения, поскольку показаны солидусы безводного плавления; в природе мокрое плавление будет понижать солидус, а источник для бонинитовых серий является более деплетированным и тугоплавким, чем источник для MORB. Tp и Tliq рассчитаны по выборке соответствующих условиям моделирования составов с помощью программы PRIMELT2.XLS [Herzberg, Asimow, 2008]. Источники данных: Троодос [Pearce, Robinson, 2010], Тонга [Falloon et al., 2008], Абитиби [Kerrich et al., 1998], Северо-Карельский пояс [Shchipansky et al., 2004]. Для сравнения показаны данные по коматиитам Косто-мукши по [Puchtel et al., 1998].

general problems of decompression melting with regard to the mantle columns formed in differing, i.e. plume or mid-ocean ridge, settings. Using the data on the most primitive compositions, that are mainly controlled by olivine fractionation, it becomes possible to calculate parental melt compositions and potential mantle temperatures corresponding to the melting initiation depth of the given mantle column [Herzberg et al., 2007; Herzberg, Asimow, 2008].

Figure 7 illustrates relationships between potential and liquidus temperatures of the most primitive compositions of the modern (Tonga forearc and Troodos and upper pillow lavas) and Archean (Abitibi and North Karelian greenstone belts) boninite series. The potential mantle temperatures for their generation differ by 150-200 °C. This difference refers to the assemblages generated by mantle sources of an almost similar type (mantle harzburgite) due to its melting in the

12

10

О

CD

Melt fraction

♦C

L+OI+Cpx+Opx+Grt

^ North Karelian

Greenstone Belt ♦ Abitibi

О Flin Flon

о Jiangnan

-^Troodos ■ Tonga

2 GPa.

L+OI+Cpx +Opx+Sp

MORB

Liquids Fertile Peridotite Accumulated Fractional Melting Dashed line = Initial melting pressure Dotted line = Final melting pressure

10

MgO (wt%)

20

30

Fig. 8. Change in parental melts of boninite series compositions in time on the petrogenic grid MgO vs. FeO [Herzberg, Asimow, 2008].

Dotted curves with arrays are fractionation paths from the primary melts to evolved boninite compositions. Data source are for Tonga [Falloon et al., 2008]; Troodos [Cameron, 1985], Jiagnan [Zhao, Asimow, 2014]; Flin Flon [Wyman, 1999a]; Abitibi [Kerrich et al., 1998]; Iringora ophiolite, North Karelian greenstone belt [Shchipansky et al., 2004]. See the text for explanations.

Рис. 8. Изменение родительских составов бонинитовых серий во времени петрогенетической сетке MgO - FeO [Herzberg, Asimow, 2008].

Точечными кривыми со стрелками показаны тренды фракционирования от первичных расплавов к бонинитовым составам. Источники данных составов: Тонга [Falloon et al., 2008]; Троодос [Cameron, 1985], Цжэцзян [Zhao, Asimow, 2014]; Флин Флон [Wyman, 1999a]; Абитиби [Kerrich et al., 1998]; Ириногорские офиолиты Северо-Карельского пояса [Shchipansky et al., 2004]. См. объяснения в тексте.

presence of hydrous fluid. So, this phenomenon may be thought of as a result of the secular cooling of the Earth. The gradient of 50-70 °C/Ga suggests evolutionary changes in geodynamic processes in the Earth's history, rather than any significant transformations in tectonic mechanisms responsible for the continental crust growth. A nearly similar difference in the potential temperatures for the plume magmatism in the Earth's history is also noteworthy [Herzberg et al., 2007].

The evolution of boninitic volcanism in the Earth's history is clearly revealed in Figure 8 with regard to the definition of potential mantle temperature, envisaging a relationship between a mantle melting degree and a depth whereat the magma column is initiated due to decompression melting. The petrogenetic grid is used to show the calculated primary melt compositions of the Late Archean (North-Karelian and Abitibi belts), Paleoproterozoic (Flin Flon belt), Neoproterozoic (Jiangnan belt) and Meso-Cenozoic (Troodos, Tonga) boninite series, and a few fractionation trends are given for illustration of differentiation paths. It is clear

that the early Precambrian boninite series were generated at higher degrees of mantle harzburgite melting (30-40 %), and the mantle melting columns occurred at considerably larger depths (3.5-4.0 GPa) than during the Phanerozoic eon (2.5-3.0 GPa). The knowledge of the P-T differences of the boninite series petrogenesis is important for understanding the evolution of geodynamic processes in the Earth's history that will be discussed below.

5. Geodynamic settings of boninite occurrence

Almost all the boninitic magmatism occurrences, regardless of its age, are genetically associated with in-tra-oceanic subduction zones (see Table 2). The modern boninitic volcanism occurrences revealed by the recent studies are localized only in the intra-oceanic plate convergence zones, and there is no proven example that would evidence any other geodynamic settings for their generation. Considerations of petrogenesis of

the boninite series rock melts also imply their emplacement above the zones wherein the oceanic lithosphere slabs are dehydrating and sinking into the mantle. It is established that in the modern intraoceanic island arcs (Izu-Bonin-Mariana, and Tonga-Kermadec), the boninite series volcanic rocks build up the perido-tite-gabbro sequences of the forearc slopes and indicate the initiation and growth of juvenile island-arcs [Pearce et al., 1992; Stern, Bloomer, 1992; Stern, 2004].

It should be noted that a genetic link between the boninite series and ophiolites was recognized well before the deep-sea drilling and dredging studies of the forearc slopes of the Tonga and Izu-Bonin-Mariana arcs [Miyashiro, 1973; Cameron et al., 1979]. Numerous discoveries of the boninite series rocks [Pearce, 2003, 2008; Stern, 2004; Metcalf, Shervais, 2008; Sklyarov et al., 2016] have given grounds to conclude that the majority of the world's ophiolites mark paleo-zones of spreading in suprasubduction environments at the ancient oceanic plate margins, rather than past spreading at ancient mid-ocean ridges. Thus, boninites are petro-logically unique rocks providing a perfect indication of paleogeodynamic settings in the oceanic plate convergence zones through the Earth's history.

The new paradigm of geodynamic settings for the boninite series generation was based on the hypothesis that SSZ ophiolites are related to the initiation of intraoceanic island arcs [Stern et al., 1991; Stern, Bloomer, 1992]. This raised the questions of how and where subduction zones were initiated [Stern, 2002, 2004]. In terms of mechanics, a prerequisite for subduction initiation is gravitational instability in the oceanic lithosphere, which can lead to its tearing followed by decompression melting of the upper mantle and sinking of one plate under another. This phenomenon, as well as the genetic link between boninites and ophiolites, is encapsulated as the subduction initiation rule (SIR) [Whattam, Stern, 2011].

In the model proposed in [Stern, 2004], this condition is satisfied when plates of differing temperature and density are juxtaposed across a transform fault or fracture zone, i.e. theoretically, in case of the interaction between plates of different ages, i.e. an old cold plate and a young hot plate. Along the fault, the gravita-tionally instability of the ancient crust makes it sink with a down-dip component of motion, while a lateral component is lacking. The model suggest that upon initiation of sinking of the slab/plate, the overlying slab/plate is subject to stretching, the partially depleted mantle is rising to melt due to its decompression, and a large input of water from the sinking slab sets up conditions for extremely high extents of fusion. Later on, the slab motion vector acquires a lateral component, which stops the decompression, and, consequently, terminates melting of the depleted mantle. Thus, a true subduction regime is established for generation of

the normal island-arc tholeiitic and calc-alkaline volcanic series.

Another subduction initiation concept is proposed in [Niu et al., 2003]. It establishes a link between locations of subduction zones and density inhomogeneities at the borders of the normal oceanic lithosphere and the oceanic lithosphere thickened by mantle plume impingement, i.e. oceanic plateaus or hotspot tracers -aseismic ridges or seamounts. This concept is supported by the data on the Tonga boninites that are confined to the intersection of the aseismic ridge of the Louisville hotspot and the paleotrough of the Tonga-Kermadec arc [Turner, Hawkesworth, 1997]. There is also evidence that the initial stage of the Izu-Bonin-Mariana arc development was associated with mantle plume fingerprinting at the Manus back-arc basin [Macpherson, Hall, 2001].

The boninite series volcanic rocks often have enriched mantle-plume isotopic and geochemical signatures [Sobolev, Danyushevsky, 1994; Taylor, Nesbitt, 1998]. The involvement of the mantle-plume thickened lithosphere in the petrogenesis of boninites would thus resolve the long debate on anomalously high temperatures for the initiation of partial melting of the depleted refractory mantle harzburgite as high heat inputs can be ensured by the ascending hot mantle plumes [Falloon, Danyushevsky, 2000].

The presence of mantle-plume products is also revealed in the SSZ ophiolite sequences, which isotopic and geochemical compositions are well studied, such as the Pindos, Josephine, Koch, Magnitorsk (Southern Urals, Russia) ophiolites, etc. (see Table 2). In the early Precambrian sequences, the boninite series volcanic rocks are also associated with mantle-plume derivatives, komatiites, as well as OIB-type metavolcanic rocks [Shchipansky, 2008].

Apparently, the occurrence of mantle-plume derivatives in the intra-oceanic subduction initiation zones does not seem to be random. It is recognized that a rising mantle plume can decrease the strength of the lithosphere, which may lead to the breakup of the continents [Courtillot et al., 1999]. In addition, an emplacement of mantle plume head at the lithosphere can significantly change its density characteristics. The ingress of melts generated by the enriched deep source into the upper layers of the mantle and the oceanic lithosphere would lead to refertilization of the previously depleted mantle. While cooling, the upper mantle transformed by mantle plume impingement becomes denser, and a new lithospheric segment may gradually become negatively buoyant. The reason is that OIB volcanic rocks are considerably enriched in Fe and Ti. Besides, Fe-Ti basalts/gabbros are known to being ec-logitizated faster than their magnesium equivalents. Figure 9 shows the experimentally determined garnet stability fields for basalt compositions differing in

2.0

1.8

1.6

1.4 —

1.2

P, G Pa

Relative proportions of Grt abandunces in the experimental runs at T=730 °C, P=1.8-2.0 GPa

Grt tholeiite basalt, high-Mg basalt

T,°C

650

700

750

800

Fig. 9. Experimentally determined Grt-in curves for com-positionally different basalts: Fe-rich basalt (Mg#=41), olivine tholeiite (Mg#=55), and high-Mg basalt (Mg#=69). Note that relative abundances of garnet shown at the inset vary significantly also. Modified after data of [Molina, Poli, 2000].

Рис. 9. Экспериментально определенные поля стабильности граната для базальтов различного состава: железистого базальта (Mg#=41), оливинового толеита (Mg#=55) и высокомагнезиального базальта (Mg#=69). Заметим, что относительные количества граната, показанные на врезке, также сильно варьируются. Модифицировано из работы [Molina, Poli, 2000].

magnesium numbers. Clearly, ferriferous compositions show the earliest occurrence of garnet in quantities considerably exceeding those in case of olivine tho-leiites and, especially, high-Mg basalts. This metamor-phic transformation at the crustal base of the thickened oceanic island builds seems to be a critical factor disturbing the gravitational stability of the lithosphere. Garnet is denser by almost 15 % than pyroxenes, amphiboles and olivine, which means that 'density eclo-gitization' can take place in the reworked lower lithosphere. In this case, even if only the lower crust is foundered, tearing of the lithosphere seems to be inevitable.

Actually, potential gravitational instability of the lithosphere seems to be directly dependent on a volume of the mantle-plume productivity and the thermal status of reworked lithosphere segment. It is not common that such factors are favorably combined; anyway, lithosphere breakup and initiation of new subduction

zones is highly probable. It is noteworthy that the polarity of subduction is predetermined by the thermal status of the lithosphere - hot lithosphere is positively buoyant, while cold lithosphere is negatively buoyant.

The model in Figure 10 shows the global-scale chemical and density inhomogeneities of the lithosphere as a result of a superplume event The term 'superplume' was coined by Roger Larson in the early 1990's, when the author has advocated that mantle-plume vol-canism was the strongest in the mid-Cretaceous time (124-83 Ma) [Larson, 1991]. This event resulted in the emergence of giant oceanic plateaus, such as the On-tong Java, Manihiki, Kerguelen, and Columbia-Caribbean. It is believed that four currently active hotspots in the Pacific Ocean, which are traceable in the oceanic seafloor, are associated with this superplume. Few other tracks seem to have been cut off in the subduction zones. Other hotspot tracks in the Pacific are younger (Paleogene-Neogene) and may be not associated with deep plumes [Clouard, Bonneville, 2001]. Accordingly the numerical simulation of a termochemical superplume at the core-mantle boundary, protuberance emissions of the deep matter to the surface are represented by hot jets or large igneous provinces which impacts are observed even at the upper mantle [Farnetani, Hofmann, 2011]. This phenomenon may be responsible for the occurrence of the first-order chemical and density inhomogeneities and the initiation of the most extended plate convergence zones. This scenario seems to be realistic for the initiation of the Izu-Bonin-Mariana arc in the early Eocene.

Another subduction initiation scenario assumes a closer link between subduction and the geodynamics of individual hotspots. As already mentioned, the youngest boninitic magmatism occurrences in the northern Tonga are spatially associated with the Louisville hotspot track. Its forearc structure includes the Vitiaz paleo-trench (age of ~4 Ma) and the modern trench located further in the ocean at a distance of about 500 km. A unique double-subduction picture is revealed in this region by detailed seismotomographic surveys. A fragment of the younger and more gently dipping slab of the oceanic lithosphere is clearly visible above the current subduction zone [Chen, Brudzinski, 2001]. The shallow-depth slab detachment was discussed in detail in [Shchipansky, 2008] as the factor disturbing a lithospheric barrier and creating extremely favorable conditions for generation of the ophiolitic boninite series (Fig. 11).

It is quite possible that the shallow-depth slab break-off resulted from the compositional density in-homogeneity of the sinking lithosphere, such as local charging by OIB magmatism products. The key geody-namic implications of this phenomenon are, first, strong short-term thermal disturbance over the narrow slab window, and, second, fast uplift of its over-

locus of future locus of future

subduction initiation subduction initiation

Fig. 10. Sketch illustrating an emergence of large refertilizited, or reworked, mantle (RM) region due to a superplume event.

This gives rise to a creation of lateral compositional buoyancy contrast over normal oceanic lithosphere (DM) and, thus, to a subduction initiation with time. Superplume cartoon is after the numerical simulation of a thermo-chemical plume [Farnetani, Hofmann, 2011].

Рис. 10. Схема, иллюстрирующая модель образования больших объемов рефертилизированной, или переработанной, мантии (RM) в результате суперплюмового события.

В результате со временем возникает химико-плотностной контраст по отношению к нормальной океанической литосфере (DM), что и определяет места инициации субдукции. Изображение суперплюма дано по результатам численного моделирования термохимического плюма [Farnetani, Hofmann, 2011].

riding plate [van de Zedde, Wortel, 2001; Buiter et al., 2002]. This mechanism provides a reasonable explanation for the short-term (3-5 Ma) volcanism which was by far more voluminous that the steady-state subduction volcanism [Stern, 2002, 2004]. Due to uplifting of

the suprasubduction plate, an ophiolite 'platform' replaced the hanging plate and provided a basement for the nascent island-arc build. It is essential that the uplift of the ophiolite 'platform' could have attained the morphological stability as the previous episode of

Low-P melting of oceanic crust

Fig. 11. Cartoon depicting a variety of the partial melting processes during shallow-level slab break-off [Shchipan-sky, 2008].

Note that in this case different portions of the upper mantle, i.e. lherzolitic asthenospheric and harzburgitic lithospheric, should be partially molten. Crustal part of the detached slab provides a source for aqueous fluids enriched by subduction-related chemical components. Loci of low pressure and high pressure melting of metabasic oceanic rocks are shown also.

Рис. 11. Принципиальная схема процессов частичного плавления в обстановке малоглубинного детачмента слэба. Модифицировано из работы [Shchipansky, 2008].

Обратим внимание, что в этом случае в область плавления включаются различные составляющие верхней мантии -лерцолитовая астеносферная мантия, гарцбургитовая лито-сферная мантия. Коровая часть оторванного слэба является поставщиком водного флюида и субдукционной компоненты. Показаны также вероятные области низкобарического и высокобарического (адакитового) плавления метабазитов океанической коры.

nascentl I oceanic lithosphère

island 4 7 boninite modified bv plumes

unmodifed oceanic lithosphère

growing

Dacnarc ™ boninite

Fig. 12. An illustration of the model of boninite formation when a reworked lithosphere caped by OIB volcanics initiates a non-stationary subduction.

The model is largely based on the available data from the modern Tonga arc [Sobolev, Danyushevsky, 1994; Turner, Hawkesworth, 1997; Chen, Brudzinski, 2001; Niu et al., 2003; Falloon et al., 2007, 2008; Resing et al., 2011]: a - lithosphere charged with Fe-rich melts gets negative buoyancy as it cooled, and starts to sink. Lithosphere tearing leads to raise asthenospheric mantle and triggers intensive melting that includes boninite series melts (see fig. 10); b - as a result of the shallow level slab break-off, a new subduction should develop and, as a consequence, a new trench occurs at a distance of a few hundred kilometers oceanward from the older trench. Detached slab fragments lead to a lithosphere extension, thus forming a tight back-arc trough characterized by occurrences of local multi-spreading centers where younger boninites can also be formed. See the text for more details.

Рис. 12. Иллюстрация модели формирования бонинитов подразумевающей инициацию нестационарной субдук-ции при участии переработанной и нагруженной OIB вулканитами литосферы.

Модель основана на многочисленных данных по современной островной дуге Тонга [Sobolev, Danyushevsky, 1994; Turner, Hawkesworth, 1997; Chen, Brudzinski, 2001; Niu et al., 2003; Falloon et al., 2007, 2008; Resing et al., 2011]. а - литосфера, отягощенная богатыми железом расплавами, по мере охлаждения приобретает отрицательную плавучесть и начинает погружаться. Разрыв литосферы приводит к подъему астеносферной мантии и вызывает интенсивное плавление, включая генерацию магм бонинитовой серии (см. рис. 10); b - в результате малоглубинного обрыва слэба возникает новая зона субдукции, новый желоб которой располагается на расстоянии в несколько сотен километров в сторону океана от предшествующего желоба. Оторванные фрагменты палеослэба провоцируют растяжение литосферы, что вызывает формирование узкого задугового трога, характеризующегося развитием мульти-спрединговых центров, где могут формироваться более молодые бониниты. См. текст для более детальной информации.

intense mantle melting leaved the mantle strongly depleted. This resulted in the emergence of a lithospheric keel capable to resist the convective instability of the surrounding mantle.

Figure 12 shows the above-described model based on the modern geodynamics of the northern termina-

tion of the Tonga arc. The top panel shows the subduction initiation and generation of boninites in the fore-arc region. The bottom panel shows the modern subduction zone and paleoslab fragments under the emerging Lau basin. In the samples dredged from the Lau basin, the presence of the subduction fluid compo-

nent is evidenced by higher contents of H2O in glasses and enrichment in large-ion lithophile (LIL) elements [Falloon et al., 1992; Danushevsky et al., 1993; Pearce et al., 1994]. The fact that compositions of these basalts significantly differ can be explained by contribution to their petrogenesis of various mantle sources, including mantle plume input. Besides, active eruptions of bo-ninites take place at present time [Resing et al., 2011]. Thus, boninites occur in a variety of combinations in different tectonic settings in time and space, i.e. in the fore- and back-arc environments. This indicates that, firstly, it is an established phenomenon, and, secondly, the subduction initiation process may be not as simple as described by the 'subduction initiation rule' (SIR) model [Whattam, Stern, 2011].

It seems likely that the subduction regime should be stabilized after a certain period of accommodation associated with the slab break-off event/ or events and resultant contrasting tectonic regimes on the surface. During this period, movements caused by the frontal subduction compression are rapidly replaced by strike-slip motion, which leads to the occurrence of a complex rift system with both hot spot volcanism and volcanic seamounts [Falloon et al., 2007,2008].

Considering the ophiolite sequences worldwide, it has been noted that, although their magmatic che-mostratigraphic progression is almost similar [Pearce, Robinson, 2010; Whattam, Stern, 2011], the relationship between boninite series and other members of ophio-lite sections are variable and may significantly differ from the ideal ophiolite sequence, known as a 'penrose ophiolite' [Sklyarov et al., 2016]. In the literature, there are numerous cases showing the lateral variability of the ophiolite sequences within the local regions and therefore suggests non-stationary settings of ophiolite-forming processes. These increasingly compel the conclusion that ophiolites are commonly associated with the subduction initiation environ-

ments, rather than with spreading in the mid-ocean ridges.

6. Subduction initiation in the Early Precambrian

As described above, generation of the early Precambrian boninite series was developed by mantle harz-burgite melting of high degrees (30-40 %), and the mantle melt columns started at greater depth (3.54.0 GPa) as comparing with those from Phanerozoic eon (2.5-3 GPa). What are the geodynamic implications of these facts?

It is well known that the early Precambrian cratons are underlain by the depleted subcontinental litho-spheric mantle (SCLM) or a lithospheric keel (root) extending into the diamond stability field. This phenomenon has never reoccurred in the later history of the Earth, and its origin has been widely discussed in the literature [e.g. Herzberg, Rudnick, 2012; Shchipansky, 2012, and references therein]. In the Phanerozoic, strong depletion of the upper mantle is associated with generation of the boninite series in the subduction initiation zones. The above overview of the global bo-ninite occurrences gives abundant evidence of bo-ninitic magmatism in the early Precambrian, and the discoveries of ancient boninites are progressively growing. Accordingly to the recently published assessment, the volumes of the Archean boninite volcanics and komatiites appear to be almost roughly similar (Fig. 13) [Furnes et al., 2014]. Discovery of the boninite series rocks with fragments of sheeted dikes and IAT-type metabasites in the oldest preserved complex Isua strongly suggests that subduction dates back from the early Earth's geological history.

A prerequisite for subduction initiation is the oceanic lithosphere breakup through its entire thickness. It follows thus from Figure 6 that, first, the early Precam-

>¡100 о

Я 75

50-

<D О)

В

о 25-I У а>

□ Komatiitic volcanism

□ Boninitic volcanism

4.0-3.0 3.0-2.5 2.5-2.0 2.0-1.5 1.5-1.0

Age intervals in billions years

1.0-0.5

0.5-0

I Fig. 13. Percentage lithology of komatiite and boninitic volcanism represented at given age intervals, relative to total of mafic and ultramafic magmatism in the Earth history, according to [Furnes et al., 2014].

I Рис. 13. Соотношения объемов коматиитового и бонинитового вулканизма в процентах от общего объема мафит-ультрамафитовых ассоциаций в истории Земли, по [Furnes et al., 2014].

Phaneroic thin-plate tectonics

LIGHT

HEAVY

graphite

diamond

andesitic crust

A boninite, OIB

graphite

diamond

Early Precambrian thick-plate tectonics

LIGHT

HEAVY

TTG crust

boninite, komatiite

I Fig. 14. Sketch illustrating distinctions between subduction initiations in modern thin-plate and Early Precambrian thick-plate tectonic settings. The dotted lines depict a conventional graphite/graphite transition relative to lithospheric mantle.

Рис. 14. Схема, иллюстрирующая различия в инициации субдукции, в обстановках современной тонкоплитовой и раннедокембрийской толстоплитовой тектоники. Точечной линией условно показано положение фазового перехода графит/алмаз относительно литосферной мантии.

brian lithosphere as a unit should have rheological properties providing for brittle or brittle-plastic deformations. In other words, such lithosphere can be considered as a rigid body capable of resisting the con-vective instability that is an attribute of plate tectonics [Sleep, 1992]. The Archaean oceanic lithosphere thickness is estimated at 85-120 km, whereas the modern lithosphere is ~60 km thick [Herzberg et al., 2010].

Second, unlike the Phanerozoic boninite series, parental melts of the early Precambrian series were generated at depths of ~120-130 km, i.e. in the diamond stability field. Given the fact that the primitive melts of the ancient boninite series clearly have the subduction influence signatures, it is reasonable to believe that deep sinking of the slabs into the early Precambrian mantle was highly possible.

Third, extensive melting of the early Precambrian upper mantle for the period of subduction initiation implies that it was subject to strong depletion through the diamond stability field. During the accommodation stage, the new sinking slab/slabs are cut off the asteno-spheric heat input thus leading to quickly cooling of overriding forearc lithosphere. This model provides an explanation of the origin and stability of the cold diamond-bearing SCLM that has not been subject to any convective perturbations, at least, since 3.0 billion years [Boyd et al., 1985].

The differences between the modern and early Pre-cambrian subduction initiation settings are schematically shown in Figure 14. The main distinction is that thin-plate tectonics covers the period from the Neopro-

terozoic to the present time, and thick-plate tectonics refers to the early Precambrian. Today, the oceanic crust is ~6-7 km thick on average, and this requires ~6-7 % melting. In general, about 1 km of oceanic crust is produced for every 1 % of partial melting (i.e. 1 km/1 % melting) [Herzberg, Rudnick, 2012]. According to available petrological estimations, the Archean MORB-type oceanic crust thickness ranges from ~20-25 km [Abbott et al., 1994] to 40-60 km [Herzberg et al., 2010]. In other words, the Archaean oceanic crust should have been, roughly, 3 to 6 times thicker than the modern crust. Such estimations scatter is due to the problem of validity of mantle potential temperature (Tp) computations for the ambient Archaean upper mantle as preserved off-arcs mafic volcanics of that age are very seldom [Pearce, 2008; and others].

Anyway, thickening of the Archean oceanic crust is a prerequisite for generation of tonalite-trondhjemite-granodiorite (TTG) gneisses composing the bulk Ar-chean continental crust. Both tonalite granitoids and their high-pressure analogues (adakites) are observed in modern suprasubduction settings, although in much smaller amounts than the basalt-andesite-rhyolite series.

7. Conclusion

The boninite series and its most fractionated end-member, boninites, are not just a simple unit of the suprasubduction sequence, but also a vivid indicator

of initiation of intra-oceanic convergence zones. Being directly related to ophiolites, boninites and low-Ti basalts/picrites give direct evidence that the oceanic lithosphere stretching was related to subduction initiation. This means that the age of SSZ ophiolites or, more precisely, SIR ophiolites [Whattam, Stern, 2011] may show the age of paleo-oceans only in the first approximation. The oceanic lithosphere itself may be much older. Indeed, the modern boninites of the northern Tonga arc, the Eocene IBM boninites and ophiolites of Papua New Guinea cannot evidently correspond to timing of the Pacific or Indo-Austarlian plate life. Moreover, the Phanerozoic folded belts show multi-su-turation signatures rather than mono-suture architecture.

The records of MOR-type ophiolites are very seldom as comparing with the large occurrences of the SIR ophiolites including the well-known sequences of Troodos, Semail, or Newfoundland [Shervais, 2001; and others]. The completely developed MOR-type includes ophiolites of 12-9 Ma age in the Macquarie Island located at the boundary between the Pacific and Australian plates, and the Neoproterozoic ophiolites of Gabal Gerf, the Arab-Nubian shield [Pearce, 2008; Whattam, Stern, 2011, and references therein]. A rarity of the MOR type in the ophiolite continuum stems well from the fact that physically it is almost impossible to emplace true MORB crust at a convergent plate boundary; rather sediments and fragments of seamounts may be scraped off of the sinking plate [Stern, 2004].

It is now recognized that SIR ophiolites are volu-metrically major in the Earth's geological history (see Table 2 and [Furnes et al., 2014]). As shown above, a prerequisite for subduction initiation is a lateral compositional buoyancy contrast within the lithosphere

which would led to its collapse and consequently to form zones of intra-oceanic plate convergence. An impingement of mantle plume into the pre-existing lithosphere appears to be the most obvious among the known mechanisms controlling the global geodynamics of the Earth. However, mantle plumes themselves do not seem to be responsible for continental crust growing even in the case of Iceland plateau placed directly above the hot mantle plume head [Martin et al., 2008]. However recently subducted oceanic plateaus crust revealed the capability to produce a juvenile continental crust similar to Archean TTG via its partial hydrous melting [Hastie et al., 2010]. It is vital also, that high-pressure adakites are also found in association with the Tonga boninites related to the hot spot track [Falloon et al., 2008]. Furthermore, the mere fact that boninites tend to be compositionally close to the bulk composition of continental crust suggesting that the crustal growth processes began with the subduction initiation (see Fig. 3).

Thus, lacking direct influence on crustal growing at the intra-oceanic island-arc zones, mantle plume events could have acted as 'remote triggers' of theses process and predetermined subsequent loci of subduction initiation zones. Such a kind of the self-organized geodynamic system of the Earth may have originated as early as Eoarchean, and its further evolution was mainly operated by the secular cooling of the planet.

8. Acknowledgements

This work is a contribution to National Task 01201459184. It was also supported by Grant 16-0500479 of the Russian Foundation for Basic Research.

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Shchipansky, Andrei A., Doctor of Geology and Mineralogy, Lead Researcher Geological Institute of RAS 7 Pyzhevsky lane, Moscow 119017, Russia И e-mail: shchipansky@mail.ru

Щипанский Андрей Анатольевич, докт. геол.-мин. наук, в.н.с. Геологический институт РАН 119017, Москва, Пыжевский пер., 7, Россия И e-mail: shchipansky@mail.ru