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Geology, tectonic settings and iron ore metallogenesis associated with subma-rine volcanism in China: An overview

Tong Hou, Zhaochong Zhang, Franco Pirajno, M. Santosh, John Encar-nacion, Junlai Liu, Zhidan Zhao, Lijian Zhang

PII: S0169-1368(13)00174-1DOI: doi: 10.1016/j.oregeorev.2013.08.007Reference: OREGEO 1080

To appear in: Ore Geology Reviews

Received date: 21 April 2013Revised date: 1 August 2013Accepted date: 8 August 2013

Please cite this article as: Hou, Tong, Zhang, Zhaochong, Pirajno, Franco, Santosh, M.,Encarnacion, John, Liu, Junlai, Zhao, Zhidan, Zhang, Lijian, Geology, tectonic settingsand iron ore metallogenesis associated with submarine volcanism in China: An overview,Ore Geology Reviews (2013), doi: 10.1016/j.oregeorev.2013.08.007

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Geology, tectonic settings and iron ore metallogenesis associated with

submarine volcanism in China: an overview

Tong Hou1, Zhaochong Zhang

1*, Franco Pirajno

2, M. Santosh

1, 3, John Encarnacion

4, Junlai Liu

1,

Zhidan Zhao1, Lijian Zhang

5

1. State Key Laboratory of Geological Process and Mineral Resources, China University of Geosciences, Beijing,

100083, China

2. Centre for Exploration Targeting, University of Western Australia, Crawley, WA, 6009, Australia

3. Division of Interdisciplinary Science, Kochi University, Kochi 780-8520, Japan

4. Department of Earth and Atmospheric Sciences, Saint Louis University, 3642 Lindell Boulevard, St. Louis, MO

63108, USA

5. No. 4 Geological Party of Hebei Bureau of Geology for Mineral Resources Exploration, Chengde, 067000,

China

* Corresponding author: Z.C. Zhang. E-mail: zczhang@cugb.edu.cn

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Abstract

Submarine volcanogenic iron oxide (SVIO) deposits are one of the most important sources of

high-grade iron ores in China. The spatial distribution of the deposits is controlled by the tectonic

settings including arc, back-arc and rifts environments, with the SVIO deposits mostly concentrated in

the western part of China namely, the southwestern Yangtze Craton, Western and Eastern Tianshan,

and Altay orogens and the Kaladawan iron ore district in the eastern part of the Altyn Tagh region. The

Chinese SVIO deposits range in age from Paleoproterozoic to Mesozoic, and were formed during two

main metallogenic epochs in the Proterozoic and Paleozoic. More than 70% of the SVIO deposits

formed in the Paleozoic, with three important SVIO–metallogenic provinces recognized, in the Altay,

Eastern and Western Tianshan orogens. These SVIO deposits are hosted in lithofacies that are related

to submarine magmatism, such as lavas and associated pyroclastic and volcaniclastic-sedimentary

rocks. The iron orebodies are hosted in different volcanic lithofacies with different features. Moreover,

the different volcanic lithofacies in which the Fe ores are hosted also provide information as to their

spatial relationship, ranging from distal to proximal to the eruption centre or vent. Many of these

deposits are characterized by well developed skarns, and could be interpreted either by a distal position

of the ore system in question and/or exposed igneous rocks or active magma chamber, or a relationship

to early metamorphism and continuous alteration at relatively high temperature followed by retrograde

alteration as temperatures decline. Geological and geochemical evidence suggests that these deposits

were formed as a result of submarine magmatic activity, including subaqueous volcanic eruptions,

associated volcano-sedimentary lithofacies, and related post-magmatic hydrothermal activity. Iron

oxide ore probably formed the hydrothermal fluids which generated the skarns could be a mixture of

evolved magma-derived water and convecting sea water driven by the heat from the shallow active

magma chamber, whereas volcano-sedimentary deposits could be formed by the fallout of the

ore-bearing materials to the sea floor emanating from submarine eruption columns, or fractional

precipitation of iron which had been introduced locally into the bottom water by volcanic-origin

hydrothermal solutions and by leaching from the relatively iron-rich volcanic rocks. The formation of

these various styles of Fe ore deposits is controlled by several key factors, such as magma

differentiation, lithofacies of host rocks, temperature and chemical compositions of hydrothermal fluids,

as well as the depth of sea water. In combination with their geological characteristics, geodynamic

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mechanisms and metallogenesis, we propose a genetic model in which the origin of these deposits can

be related to the space-time evolution of the submarine volcanism, and their relationship to volcanic

lithofacies variation, such as central, proximal and distal environments of ore formation.

Keywords: metallogenesis; iron deposits; submarine volcanic rocks; China

1. Introduction

Three-quarters of the modern Earth’s volcanic activity is submarine, located predominantly along

the mid-ocean ridges, with the remainder along intra-oceanic volcanic arcs and hotspots, at sea floor

depths varying from greater than 4000 m to near the sea level (e.g., Carey and Sigurdsson, 2007;

Embley et al., 2006). Submarine volcanic eruption is difficult to observe directly, and their products

are difficult to recover and study. Hence, evidence of submarine volcanism comes from sightings of

explosive sea level manifestations (Kokelaar and Busby, 1992).

It is widely recognized that these volcanoes play a role in transferring mass and energy from the

oceanic crust and mantle to the oceans, which is a favorable environment to form metal-rich deposits

(e.g. Tivey, 2007) as demonstrated by the abundant Fe and base metal deposits present on land formed

during geologic history, such as Algoma-type BIF and VMS deposits (Mücke et al., 1996; de Ronde et

al., 2005). The metallogenesis of these deposits, and the distribution and composition of submarine

volcanic systems that create them had been relatively well studied. In contrast, many, and probably

most, iron oxide deposits associated with submarine eruptions, especially those generated in the

Phanerozoic have not been investigated in detail yet.

In China, the discovery of many iron oxide deposits associated with submarine volcanic rocks is

considered as one of the last century’s most exciting facets of geological research on iron oxide ore

deposits (e.g. Jiang and Wang, 2005). Submarine volcanogenic iron oxide (SVIO) ore systems mainly

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include volcanic–associated and (volcano sedimentary)–hosted. The iron oxide ores typically occur as

lenses, layers and veins that may form at or near the seafloor in submarine volcanic environments.

They have been regarded to be formed by iron-enriched melts/fluids associated with seafloor volcanic

eruptions, linked to submarine hydrothermal systems (Kelly, 2002, Hannington 2005; see Pirajno 2009

for an overview). SVIO deposits in China are possibly related to a wide range of geodynamic settings

and depositional environments, such as island-arcs, rifts and mid-ocean ridges and oceanic islands. It is

noteworthy that most of the SVIO of China are composed predominantly of high-grade iron oxide ores,

thereby contributing a considerable amount of iron for the local industry (Jiang and Wang, 2005; see

also Hu et al., 2011).

However, although these SVIO deposits have attracted a substantial number of petrologic and

geochemical studies (e.g., Jiang and Wang, 2005), their metallogenesis and the genetic relationship

with associated submarine volcanism are still poorly understood, with various genetic models proposed,

including sea floor volcanic systems, skarn and exhalative-sedimentary (e.g., Feng et al., 2009; Hua,

1985; Shan et al., 2009; Zhang et al., 1987).

The previous studies of SVIO deposits of China have shown some similarities as well as differences

from their subaerial counterparts (e.g. Wang and Chen, 2001; Jiang, 1983). For example, ores formed

by eruption of iron oxide melt can be compared with the Kiruna style mineralization, such as the El

Laco deposit in Chile (Henríquez et al., 2003). On the other hand, leaching of ore-bearing pyroclastics

by deep sea water as one of the major sources of iron for the SVIO deposits is seldom seen in terrestrial

environments. In this paper we present an overview of the geological characteristics, and geodynamic

mechanisms of the Chinese SVIO deposits, comparing them with the actively forming iron deposits

along modern subduction zones, mid-ocean ridges, and back-arc basin in order to refine our

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understanding of the metallogenesis of SVIO deposits. Furthermore we also provide a comprehensive

overview based on published works on submarine volcanic processes and the related iron oxide

deposits. At the end of the paper, we propose a genetic model which links the origin of these deposits to

the space-time evolution of the submarine volcanoes, and integrated them on the basis of principal

volcanic lithofacies variation according to their closeness to vent, i.e. central, proximal and distal

facies.

It is worthwhile to point out that, in spite of the ancient Algoma-type BIFs being closely related to

submarine volcanism (Mücke et al., 1996), in most cases they have been subjected to varying degrees

of alteration, deformation, and metamorphism resulting in the destruction of the original textures and

structures. The origin of the Algoma BIF deposits is therefore beyond the scope of this paper and will

not be addressed.

2. Distribution of SVIO deposits and geological setting

The tectonic framework of China is dominated by three major Precambrian cratons, the North China,

South China (Yangtze+Cathysia) and Tarim Cratons (Fig. 1), surrounded by fold belts and accretionary

orogens including accreted island arcs, back-arcs and oceanic lithosphere (Zhai and Santosh, 2011,

2013).

Submarine volcanogenic iron oxide (SVIO) deposits, one of the most important iron deposit types in

China, have been recognized to be widely distributed in volcanic provinces, mostly located in western

China. These deposits cover a considerable age range, from Proterozoic to Mesozoic, but with more

than 70% of SVIO deposits formed in the Paleozoic, especially in Late Paleozoic. Several important

SVIO-metallogenic provinces have been recognized in the Western Tianshan, Eastern Tianshan, Altay,

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Kaladawan area at eastern part of the Altyn Tagh Mountain and western margin of Yangtze Block

(Fig.1). The geological settings of the main SVIO-metallogenic provinces are summarized in the

sections that follow.

2.1. Western Tianshan

The Chinese Western Tianshan Mountain is located along the southwestern margin of the Central

Asia Orogenic belt (CAOB), and represents a Neo-Proterozoic-Paleozoic orogenic belt extending from

the Siberian Craton in the north to the Tarim Craton in the south (Xiao et al., 2004; Windley et al.,

2007; Wong et al., 2010; Xiao et al., 2013; Rojas-Agramonte et al., 2011; Kröner et al., 2007). The

Chinese Western Tianshan Mountain is a late Paleozoic accretionary orogenic belt (Fig.2; Allen et al.

1992; Gao et al. 1998, 2009) where the passive margin of the northern Tarim plate finally amalgamated

with the active margin along the southern Siberia plate. The Late Paleozoic tectonic evolution of the

Chinese Western Tianshan Mountain can be broadly subdivided into two stages (Gao et al., 1998; Chen

et al., 1999; Gao and Klemd, 2003): 1) dominantly subduction, expressed by the southward subduction

of the North Tianshan Ocean (e.g. Wang et al., 2008) or northward subduction of the South Tianshan

Ocean (Gao et al., 1998, Long et al., 2008) beneath the Yili block, and north-directed A-type

subduction of the Tarim Plate, followed by exhumation; 2) dominantly a transition from subduction to

post-collisional extension at ca. 320 Ma (Gao et al., 2009; Sun et al., 2008).

The exposed strata include Proterozoic, Silurian, Devonian, Carboniferous, Permian, Triassic,

Jurassic and Quaternary (e.g. Sun et al., 2008). Among of above, the Carboniferous and Silurian rocks

are most widely distributed. Magmatism, both intrusive and extrusive took place throughout the Early

Paleozoic and Late Paleozoic. Early Carboniferous (Mississippian) and Early Permian volcanic rocks

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are also well developed. Most of the igneous rocks are intermediate-felsic, or intermediate-mafic. High

grade SVIO deposits, most of which were discovered in the Awulale Metallogenic Belt, are spatially

and temporally associated with the submarine volcanic rocks of Mississippian Dahalajunshan

Formation. The tectonic setting for these volcanic rocks is still controversial with two contrasting

models currently proposed for the Dahalajunshan Formation, 1) extensional setting (e.g., Che et al.,

1996; Xia et al., 2004) and 2) Late Paleozoic continental arc setting related to the southward movement

of the North Tianshan Ocean (Wang et al., 2008) or northward subduction of South Tianshan Ocean

(Gao et al., 1998, Long et al., 2008) beneath the Yili block, respectively. Based on recently published

age data, more researchers favor the second model, which relates the formation of the volcanic host

rocks of the submarine iron ore deposits in the Western Tianshan to a late Paleozoic subduction process

(Zhang et al., 2012). More detailed descriptions of the iron deposits associated with submarine volcanic

rocks in the Chinese Western Tianshan Mountain can be found in Zhang et al. (this volume).

2.2 Eastern Tianshan

The SVIO in the Eastern Tianshan forms a belt, located between the Junggar block and Tarim block

(Fig.3). The Paleozoic tectonic evolution history of the Eastern Tianshan remains controversial. Some

researchers have suggested that the Eastern Tianshan results from the southward subduction of the

Junggar Ocean along the Bogda-Haerlike zone (Qin, 2000; Zhang et al., 2004), while others have

proposed a northward subduction of the south Tianshan ocean instead (Wang et al., 2006).

The Eastern Tianshan area is bound to the north by the Turpan-Hami (usually abbreviated to Tuha)

basin, which is a part of the Junggar block, and to the south by the Aqikekuduke fault, which separates

this northern belt of the Tianshan from the Central Tianshan.

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Geological mapping and geochemical surveys, have identified three main tectonic domains, in the

Eastern Tianshan: 1) Dananhu-Tousuquan arc in the north (north belt), 2) Jueluotage ductile shear zone

in the middle, including the Kushui-Gandun back-arc basin border facies (middle belt) and 3) Yamansu

back-arc basin (south belt), and Central Tianshan Microblock in the south (Fig.3; Qin et al., 2002,

2003). The Jueluotage belt can be further subdivided, from north to south, into

Wutongwozi–Xiaorequanzi intra-arc basin, Dananhu–Tousuquan island arc, Kangguer–Huangshan

ductile shear zone and Yamansu (Kumutag-Shaquanzi) back-arc basin (Qin et al., 2002). The

Bogda–Haerlike belt is made up of well developed Ordovician–Carboniferous volcanic rocks intruded

by late Paleozoic granites and mafic–ultramafic complexes (Gu et al., 2001; Li et al., 2006; Ma et al.,

2013). The typical high grade iron ore deposits associated with submarine volcanic rocks, such as

Yamansu, Kumutag, Bailingshan and Hongyuntan deposits are found in the Yamansu back-arc basin

(Fig. 3).

The Yamansu back-arc basin lies between the Aqishan-Yamansu fault (or Kushui fault), which

marks the southern boundary of the Kanggurtag shear zone, and the Aqikekuduke fault. The exposed

rocks comprises a 5 km thick succession of Lower Carboniferous Yamansu Formation bimodal

volcanic rocks, middle Carboniferous flysch of the Shaquanzi Formation, and Upper Carboniferous

clastic rocks, andesitic tuff, and intercalated carbonate rocks of the Tugutublak Formation. The

Carboniferous rocks are overlain by the Permian marine and terrestrial clastic rocks which are

intercalated with bimodal volcanic rocks and carbonate rocks. Carboniferous–Permian magmatism was

extensive and resulted in the emplacement of high-Na, relatively oxidized, calc-alkaline to alkali

magmas (Qin et al., 2002, 2003).

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2.3 Altay

The ~500 km-long Altay orogenic belt in NW China (Xinjiang Uygur Autonomous Region) is

separated to the southwest from the Junggar terranes by the Erqis (also known as Ertix, Irtysh)

strike-slip fault (Fig.4; Yu et al., 1993; Qin and Dong, 1994). The geodynamic evolution of the Altay

orogeny remains controversial. Felsic magmatism and translithospheric strike-slip movements suggest

that the collision of the Siberian Plate and Kazakhstan Block (Junggar Terrane) occurred between the

Early and Late Carboniferous, resulting in the accretion of island arcs and other terranes, which

constitute the Altay orogenic belt (Li and Poliyangsiji, 2001; Li and Zhao, 2002; Xu et al., 2003; Yang

et al., 2007). Paleontological and paleomagnetic studies argue for an Early Permian collision (Cocks

and Torsvik, 2007). Xiao et al. (2008) proposed that the formation of the complex orogenic collage

between the Siberian Plate and Kazakhstan Block occurred between Late Permian and Triassic times.

Nevertheless, more recent studies suggest Late Silurian to Early Devonian magmatism at the southern

margin of the Chinese part of the CAOB occurred in an active continental margin setting (c.f. Chai et

al., 2009).

The Early Paleozoic–Late Paleozoic Altay orogeny in NW China is further subdivided into the North

Altay, Central Altay and South Altay (Xiao et al., 1992; Ye et al., 1997; Yang et al., 2007). The South

Altay is characterized by Middle-Ordovician low-grade metamorphosed rocks (Habahe Group), Late

Silurian to Early Devonian Kangbutiebao Formation containing submarine volcanic and sedimentary

rocks of low grade metamorphism, and the Middle Devonian Altay Formation, consisting of

sedimentary rocks intercalated with low-grade meta-volcanic rocks. In addition, Ordovician volcanic

rocks and sedimentary clastic rocks, Silurian Kulumuti Group crystalline schists and migmatites are

also present but less commonly exposed in the area. Voluminous Early and Late Paleozoic

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syn-orogenic and post-orogenic granitoids (Tong et al., 2005; Yuan et al., 2007) and Cambrian to

Permian volcanic rocks are recognized in the Altay (Windley et al., 2002). The rocks of the

Kangbutiebao Formation are distributed in the Chonghu'er, Kelang and Maizi volcano-sedimentary

basins, all have undergone regional greenschist to lower amphibolite facies metamorphism. The Fe ore

deposits in the Kangbutiebao Formation include the Mengku and Abagong Fe and Fe-P deposits, which

consist of mafic to silicic volcanic rocks and metasedimentary rocks. Specifically, the Mengku deposit

is hosted in the lower part of Kangbutiebao Formation, whereas the Abagong deposit is hosted in the

upper part.

2.4 Kaladawan area

The Kaladawan area in the eastern part of the Altyn Tagh Mountain, which is situated between the

Tarim Basin and the Qaidam Basin in northwestern China (Fig.5; Guo et al., 1999), is located between

the NE-trending Altyn Tagh strike-slip fault and the E-W-trending Northern Altyn Tagh fault. The

tectonic evolution of Altyn Tagh Mountain is still debated (Yin et al., 1999; Sobel and Arnaud, 1999).

The Altyn Tagh has an Archean-Paleoproterozoic basement (Cui et al., 1999), overlain by Middle

Proterozoic rocks, later affected by Neoproterozoic-early Paleozoic within plate extension (Guo et al.,

1999), followed by Early Paleozoic subduction (Xu et al., 1999; Sobel and Arnaud, 1999), late

Paleozoic rift extension, orogeny and related. Triassic extension with emplacement of alkali intrusions

(Yin et al., 1999), as well as sinistral strike-slip movement occurred due to the far-field effect of the

India-Eurasian collision during the Cretaceous in Altyn Tagh fault belt, which exerted a regional

compressive regime in the Kaladawan area (Guo et al., 1999; Cui, et al., 2002; Chen et al., 2002)

In the Kaladawan area, the basement consists of Archaean high-grade metamorphic rocks such as the

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Dagelagebulake Formation, including granulite, gneisses, amphibolites and migmatites. Cambrian

meta-volcanic rocks constitute the major exposure in the area, and are overlain by Upper Carboniferous

clastic and carbonate sequences. Additionally, the Upper Oligocene Ganchaigou Formation, Middle

Oligocene Youshashan Formation and Quaternary sediments (Fig.5) occur in the southern part. The

Cambrian rocks comprise the Zhuabulake Formation and the overlying Simierbulake Formation The

former covers most of the area, consisting of dark mudstone, carbonaceous phyllite, siltstone, light grey

slate, mica schist, quartz schist, and marble, meta-dacitic rocks, felsic tuff and basalt. The basalt is the

major host rock of iron mineralization, is interbedded with sedimentary rocks, and exhibits aphanitic

texture and pillow structures. In the middle-northern part of the area, the major rock types are

sericite-schist, sericite-quartz-schist, phyllite, slate and intermediate-felsic volcanic rocks interbedded

with marble and quartzite units. The Upper Carboniferous Yingebulake Formation is made up of

sandstone, siltstone, limestone and shale, is locally exposed. Siltstone, mudstone and conglomerate

constitute the Oligocene succession, which occur only in the southeastern part of the area. The

Cambrian and Carboniferous sequences are intruded by gabbro, diorite, granodiorite and granitic

porphyry (Chen et al., 2009). In recent years, a number of SVIO deposits and occurrences have been

identified in this area thanks to high-resolution aeromagnetic surveys (Chen et al., 2009). These

deposits are exclusively found in the Early Paleozoic volcanic rocks, and it has been suggested that the

volcanic rocks formed in an early Paleozoic arc setting (Cui et al., 2010).

2.5 Southwest margin of Yangtze Block

South China comprises the Yangtze Block to the northwest and the Cathaysia Block to the southeast,

which were amalgamated along a Neoproterozoic collisional belt (Fig.6; Chen et al., 1991; Li and

McCulloch, 1996; Zhang and Zheng, 2013). To the north, the Yangtze Block is separated from the North

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China Block by the Qinling-Dabie orogenic belt, which was formed by the closure of the easternmost

part of Paleotethys in the Triassic (Mattauer et al., 1985; Wu and Zheng, 2013). To the west, it is bound

by the Tibetan Plateau. In the western part of the Yangtze Block, Mesoproterozoic granitic gneisses and

metasedimentary rocks are intruded by Neoproterozoic (~800 Ma) arc-related granites (Zhou et al. 2002b)

and overlain by a series of Neoproterozoic (~600 Ma) to Permian marine and terrestrial rocks. During the

Cenozoic, the western part of the Yangtze Block was subjected to strike–slip faulting and thrusting, while

the eastern part was dominated by block faulting and shallow-level shearing, e.g. Cenozoic

Ailaoshan-Red River Shear Zone (Burchfiel et al. 2008).

The Neoproterozoic tectonic evolution of South China has long been a matter of debate. Some

workers suggested that the Neoproterozoic (ca. 825 Ma) magmatism in South China was produced by a

mantle plume that heralded the pre-breakup of Rodinia (Li et al., 1995, 1999). On the other hand, Zhou et

al. (2002a, b) argued that the Neoproterozoic igneous assemblages along the western margin of the

Yangtze Block represent part of a magmatic arc, suggesting the presence of a major subduction zone

during the Neoproterozoic. The dominant mineral deposits associated with Proterozoic rocks in the

southwestern margin of Yangtze Block are precious and base metal (Fig. 6).

The Early-Middle Proterozoic Dahongshan Group comprises limestone, sandstone, basalt and

pyroclastic rocks, in which the volcanic units represent an Early-Middle Proterozoic volcanic activity

along the western margin of Yangtze Block. The age of this volcanism is about 1700 Ma (Rb-Sr isochron)

and 1900 Ma (single zircon U-Pb) (Hu et al., 1991; Greentree and Li, 2008). The Dahongshan Group

conformably overlies the basement, with the Archean Dibadu Formation dominated by

basaltic-andesitic volcanics with a limestone-sandstone sequence. Because of the small outcrops, and the

metamorphic overprinting, it is difficult to identify the original depositional structures. Several studies

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have demonstrated that Paleoproterozoic subduction occurred in the western margin of Yangtze Block,

and the basaltic-andesite volcanics were formed during paleo-Qinghai-Tibet oceanic plate subduction

under the Yangtze plate (e.g. Zhang et al., 2001).

3. Geology of Chinese SVIO deposits

Eruptions on the seafloor and submarine magmatism constitute by far the largest proportion of the

Earth’s volcanism. Submarine volcanic eruptions occur at divergent plate boundaries (e.g. Buck et al.,

1998; Macdonald, 1998; Perfit and Chadwick, 1998; Head et al., 1996) and intraplate regions,

commonly building seamounts (e.g. Keating et al., 1987; Wessell and Lyons, 1997; Schmidt and

Schmincke, 2000). Similar to their subaerial counterparts, central-type submarine volcanism can

produce not only multiple facies, such as lavas, pyroclastic rocks, volcano-sedimentary (volcaniclastics)

rocks, but also show similar spatial distribution of the volcanic products around an eruptive centre.

SVIO deposits have been identified to be associated with different facies, and as such they tend to have

complex geological characteristics. Therefore, detailed investigations of the SVIO deposits of China

are critical to unravel the complexities of both individual and regional-scale metallogenic processes. In

this paper, we classify the SVIO deposits in China into groups according to the types of host rocks,

namely: lavas, pyroclastic rocks, volcano-sedimentary rocks, and ore systems of uncertain or

polygenetic origin. Typical examples are briefly described in the following sections.

3.1 Submarine lava-hosted type

Geologic relationships suggest that this type of SVIO deposits contribute most reserve of SVIO ores,

and usually occurs as intercalated layers or lenses within submarine volcanic rocks, with or without

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significant occurrence of skarn minerals. Particularly, the submarine lava-hosted iron ores have been

identified at Dahongshan in Yunnan, which probably because of the role played by

volcano-sedimentary processes in the metallogenesis at Dahongshan, we will discuss it separately in

the section on uncertain or polygenetic SVIO deposits. Other typical lava hosted type SVIO deposits

include Yamansu in Eastern Tianshan, as well as the several iron deposits in Kaladawan area and

Chagangnuoer and Zhibo in the Western Tainshan. Below, we summarize the salient features of the

Yamansu deposit.

3.1.1 Yamansu

The Yamansu Fe-Cu deposit in Eastern Tianshan contains a reserve of 32 Mt with an average grade

of 51 wt.% Fe, and 20,000 t with a mean of 0.06 wt.% Cu (Mao et al., 2005). The Yamansu iron deposit

occurs about 80 km south of Hami City. Regionally, the exposed strata consist of Lower Carboniferous

Yamansu Formation, Upper Carboniferous Shaquanzi Formation, and Lower Permian Aqikebulake

Formation. Around the Yamansu open pit, the Yamansu Formation comprises intermediate-basic lava

and pyroclastic rocks, limestone and minor felsic rocks (Fig. 8). The Shaquanzi Formation mainly

comprises flysch, and is overlain by the Lower Permian marine and terrestrial clastic rocks, which are

intercalated with bimodal volcanic rocks and carbonate rocks. A number of faults have been recognized

surrounding the deposit, and they include five NNE to ENE-trending faults (Fig. 8a). The lava flows

are predominantly basaltic with minor andesite in the Yamansu deposit. The basaltic and andesitic lavas

display a gradational contact, and the two rock types cannot be easily distinguished in hand specimen.

These flows are generally several meters thick, rarely up to 100 m. The lava flows are interbedded

within pyroclastic rocks. No intrusions have been identified at the Yamansu deposit, except for the

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subvolcanic pyroxene-diorite porphyry exposed about 500 m southwest of the orebodies (Fig. 8a).

However, a gravity survey suggests that some buried intrusive rocks might be present at depth (Mao et

al., 2005). Several ancient volcanic edifices were recognized adjacent to the deposit on the basis of

remote sensing and facies analysis of the volcano-sedimentary rocks. The Yamansu volcanic lavas are

considered to be located within or adjacent to the volcanic center (BGMRXUAR, 2010).

Eighteen orebodies have been recognized in the deposit and occur as EW-trending stratiform, banded

podiform to lenticular bodies (Fig. 8). Nos. 1, 2, 4, 7 and 8 orebodies are the largest, and Nos. 1 and 2

orebodies are the most economic. No. 1 orebody is >940 m long, and dips southwards with the dip

angle of 43° at surface (980 m above sea level) to 72° at 420 m above sea level. The average width of

the No. 1 orebody is 8.6 m. The No. 2 orebody strikes ~1300 m discontinuously, dip southwards at 59°

and is 7–17 m wide. Country rocks to orebodies are mainly mafic lavas and pyroclastic rocks

intercalated with limestone of Yamansu Formation (Fig. 8b). The orebodies are mostly conformable

with their country rocks (Fig. 8b). Based on mineral assemblages, three types of ores have been

identified: garnet–magnetite, garnet–magnetite–pyrite and magnetite–pyrite (BGMRXUAR, 2010).

Field evidence and petrographic observation indicate four stages of mineralization: (1) prograde stage:

garnet + albite + apatite, (2) retrograde stage: magnetite + epidote + chlorite + quartz + amphibole +

apatite, (3) sulfide stage: pyrite +chalcopyrite + pyrrhotite + chlorite + quartz + calcite + galena +

sphalerite, and (4) supergene stage: hematite + malachite + siderite + quartz + calcite (BGMRXUAR,

2010). Magnetite is the predominant ore mineral which occurs together with minor hematite, pyrite and

chalcopyrite. The gangue minerals consist of garnet, hornblende, biotite, chlorite, epidote, quartz,

calcite and other calc-silicate minerals (Mao et al., 2005). Ore textures include massive, banded,

disseminated and irregular. The sulfide stage is dominated by pyrite, chalcopyrite and pyrrhotite. Pyrite

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occurs as cubes in massive veins (~5 mm) or as isolated grains with amphibole and plagioclase, which

often display cavities and embayed margins. Calcite and minor quartz are the main gangue minerals in

this stage. They usually cut the earlier formed minerals like garnet and amphibole as veins or

stockworks. Hematite, siderite and malachite are restricted to the supergene stage.

Skarn is ubiquitous and intensively developed in the Yamansu deposit, with a strike length of ~1000

m, a depth in excess of 600 m and an average width of 120 m as demarcated from surface mapping and

diamond drilling (Fig. 8b). The skarn shows a distinct boundary with the country rocks. The dominant

skarn minerals are garnet with subordinate amphibole, epidote, chlorite, pyroxene, albite, as well as

magnetite, pyrite, chalcopyrite and pyrrhotite. The prograde stage is characterized by formation of a

large amount of garnet. In contrast, pyroxene is very limited (~5%) and typically occurs as random

pods (Ding, 1990). The retrograde stage is characterized by hydrous alteration, and dominated by

epidote, and minor amphibole and chlorite, which replace the prograde minerals to variable degrees.

The epidote is closely associated with the magnetite (BGMRXUAR, 2010). The amphibole veins

commonly cut across garnets, indicating that amphibole formed later than garnet. During the late

retrograde stage, a large quantity of magnetite, and epidote, amphibole, chlorite and garnet formed.

Epidote is the most common mineral in the strongly retrograde altered rocks. Field relations and

petrographic studies on the mineral paragenesis reveal that the skarn at Yamansu is similar to other

conventional iron-bearing skarn deposits (Einaudi, 1981).

Whole rock K-Ar ages have a range of between 360 and 190 Ma, whereas a Rb-Sr isochron age of

286 Ma was obtained from mineralized quartz veins from a similar skarn deposit (Bailingshan), also in

the Aqishan-Yamansu rift belt (Mao et al., 2005). Recently, Hou et al. (2013) conducted laser ablation

inductively coupled plasmamass spectrometry (LAICP-MS) U–Pb zircon dating of the basalts and

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skarns yields almost coeval ages of 324.4 ± 0.94 and 323.47 ± 0.95 Ma, respectively.

3.2 Volcano-sedimentary rocks hosted

The iron oxide mineralization of this type shows a strata-bound characteristic, locally occurring in

specific beds. The ores display dominantly fine grained texture and banded and laminar structures,

which are features common for sedimentary deposits. Based on their proximity to the source (vent area),

these deposits can be classified into two sub-types. In the first type, the mineralization is located near

the eruptive centers of submarine volcanics and is defined by volcanic domes or coarse-grained

pyroclastic breccias, tuff and lava. Quartz, sericite, and chlorite alteration is common adjacent to or

beneath the deposits, indicating a possible paragenetic relationship with iron oxide (Lowman and

Bloxam, 1981). This type had been described in the last section. The second category of orebodies is

distributed in the peripheral zone of the volcanic center, and exhibits layered or stratiform shape. They

are exclusively hosted by pyroclastic rocks (e.g. tuffaceous rocks) or sedimentary rocks, such as

sandstone, dolomite and limestone, and chert. The Abagong deposit in Altay and Songhu deposit in

Western Tianshan are typical examples for this type.

3.2.1 Abagong iron

The Abagong high grade iron deposit (44.18wt.%-67.21wt.%), with accompanying P2O5

(3.8wt.%-10.8wt.%) mineralization, is located in late Silurian-early Devonian felsic volcanics along the

southern margin of Altay, Xinjiang Uygur Autonomous Region. The mineralization occurs as

structurally-controlled lenses, veins and stratiform bodies (Li and Chen, 2004; Fig.9). The iron ores are

predominantly hosted in the Kangbutiebao Formation which, as mentioned previously, comprises

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volcanic and pyroclastic rocks intercalated with sedimentary rocks, metamorphosed to greenschist and

up to amphibolite facies after emplacement. The mafic volcanic rocks of Kangbutiebao Formation have

a tholeiitic composition (Fig. 7b), but the main host rocks are metarhyolites, typically with a polygonal

granoblastic texture (felsitic, high-temperature static recrystallisation), but locally overprinted by

regional planar fabrics, probably associated with multistage strike-slip movements of the Abagong

Fault (c.f. Pirajno et al., 2011)

The length of ore bodies ranges from 200-1800 m, with thickness in the range of 1.4-16.5 m (Pirajno

et al., 2011). Magnetite dominates the proportion of ore minerals, coexisting with considerable amounts

of apatite, fluorite and lesser pyrite. Wall rock alteration minerals include tremolite, actinolite, chlorite,

albite, kaolinite, quartz, phlogopite, epidote and calcite. All of these minerals are also present in the

Fe-P ores of the Kiruna district, where they were considered part of a skarn association (e.g. Nyström

and Henríquez, 1994). However, at Abagong no skarn was noted. The nature of the Abagong

mineralisation is poorly known, with only conference abstracts, specifically addressing this deposit (e.

g. Liu et al., 2009a, b) or simply reporting the associated lithologies (e.g., Li and Chen, 2004; Chai et

al., 2009; see Pirajno et al. 2011 for an overview on Abagong Fe-P deposit). From these authors it can

be surmised that the Abagong mineralisation occurs primarily as structurally-controlled lenses and

veins. Liu et al. (2009a), on the basis of REE composition (LREE-enriched, marked negative Eu

anomalies) of the apatites as well as the magnetite-apatite ore association, classified Abagong as a

Kiruna-style mineral system. The host rocks of the Kangbutiebao Formation have been studied in some

detail by Chai et al. (2009), who performed SHRIMP U-Pb analyses of zircons from the metarhyolites,

yielding ages ranging from 412.6±3.5 Ma to 406.7±4.3 Ma. One important conclusion reached by Chai

et al. (2009) is that the magnetite-apatite ores postdate the rocks of the Kangbutiebao Formation, which

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they suggested may have formed in a subduction-related setting. The Early Devonian silicic

magmatism, now represented by metarhyolites, would have been formed by partial melting of

continental crust, whereas mafic rocks resulted from a heat source related to mafic underplating, which

then caused partial melting of the overlying continental crust. However, till now, no explanation was

offered for the magnetite-apatite ores, except that these resemble Kiruna-style mineral systems. We

admit that the Kiruna-type label is probably correct, but it must be borne in mind that the origin of

Kiruna-type Fe-P ores is controversial, although a magmatic origin is perhaps undisputed, but details

have remained conjectural since their first discovery in Sweden, some 300 years ago (Pirajno et al.

2011). Thus, the door to the Abagong Fe-P mineralization remains open and further work is needed to

unravel its origin and ore system classification.

3.2.2 Songhu

The Songhu iron deposit is located at the eastern part of the Awulale Metallogenic Belt. More

specifically the deposit is within the Yili micro-block of Kazakhstan plate, and belongs to the

Awulale-Yisjilick late Paleozoic rift system. Rocks exposed in the mining area include middle-upper

Devonian Kansu Formation, Carboniferous Awulale Formation and Tuergong River Formation,

Middle-Lower Jurassic Kashi River Formation, and Quaternary sediments (Fig.10). The Kansu

Formation consists predominantly of tuff, tuffaceous siltstone intercalated with few limestone, and

dacite. The Awulale Formation can be divided into three members. The first member comprises felsic

volcanic and pyroclastic rocks, such as volcaniclastic and rhyolitic rocks; the member is composed of

limestone, silty mudstone, and sodic rhyolite and associated volcaniclastics. The lower part of the third

member is limestone and the upper part consists of andesitic pyroclastic rocks. Iron oxide ores have

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been recognized in these rocks (Fig.10) and exhibit sharp contacts with the wall rocks. The Tuergong

River Formation is locally exposed, and consists of tuffaceous conglomerate, intermediate-felsic tuff.

The Kashi River Formation unconformably overlies the Awulale Formation and consists mainly of

conglomerate, sandy conglomerate and sandstone.

The structurally-controlled Songhu iron ore deposit, is hosted by the pyroclastic rocks intercalated

with carbonate, and is located in the northern limb of the Gongnaisi syncline. The orebodies are layered

or lensoid in shape, conformable with the host rocks. Ore types mainly consist of massive and

disseminated. Mineral assemblages comprise magnetite and hematite with subordinate amounts of

pyrite and chalcopyrite, and gangue minerals are predominantly composed of tremolite, actinolite,

epidotite, chlorite, garnet, quartz and calcite. Wall rock alteration minerals include epidote, chlorite,

carbonate and lesser magnetite and pyrite (Shan et al. 2009).

3.3 Uncertain or polygenetic iron ore systems

This type of iron ores show more variable characteristics compared to those discussed above.

Particularly, apart from the submarine volcanic activities, other processes including

volcano-sedimentary, post-magmatic hydrothermal activity etc. probably played important roles

during the iron mineralization. Notably, the large scale mineralization in these complex deposits

dominantly occur in close proximity to submarine volcanic center or directly above the volcanic vent.

The Mengku iron ore deposit in Altay, Dahongshan iron-copper deposit in Southwestern margin, and

Kaladawan iron ore district are probably typical examples for this type. At least two distinct

submarine volcanic processes were involved in the formation of iron ores. A brief description of the

salient features of some of these deposits is given in the following sections.

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3.3.1 Mengku

The Mengku Fe deposit had resources estimated at about 110 Mt of ore with grades ranging from

24 to 57.6 wt.% (Wang et al., 2003), but more recent data indicate a total resource of 200 Mt (Yang et

al., 2010), with one orebody (No. 1) containing 35 Mt, grading 41wt.% Fe (Xu et al., 2010). Wang et

al. (2003), Yang et al. (2010) and Xu et al. (2010) reported on this deposit and reviewed in Pirajno et

al. (2011).

The Mengku deposit is on the northwestern limb of anticline of Upper Silurian rocks of the

Kulumuti Group and Lower Devonian rocks of the Kangbutiebao Formation (Fig.11). The Kulumuti

Group is 6000 m thick and comprises metasandstone, phyllite, slate, biotite schist, two-mica schist,

gneiss and migmatite; the Kangbutiebao Formation totals 1300 m in thickness and comprises brown

marble, banded inpure marble (Lower Member), hornblende granulite, leptite, hornblende gneiss,

amhibolite (Middle Member) and the main host of the Fe ore (Upper Member), which consists of

hornblende-biotite-quartz schist, marble, hornblende-albite granulite and hornblende gneiss. Also in

the host sequence is a Na-rich metarhyolite (Xu et al., 2010). The upper units of the Kangbutiebao

Formation consist of a 700 m-thick sequence of metasandstone, biotite-quartz schist,

hornblende-garnet schist, marble and felsic metavolcanic rocks. Granitic rocks of assumed Late

Paleozoic age are exposed in the deposit area, comprising gneissic granite, alkali-feldspar granite,

biotite granite, two-mica granite and quartz diorite. One of the local granites is the Mengku pluton,

with U-Pb zircon ages of ca 404 to 400 Ma (Yang et al., 2010).

The Mengku deposit comprise twenty nine orebodies, ranging in shape from podiform to lenticular

to irregular and striking 120°-110°. The Fe ore is arranged in a synclinal structure within the

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northeastern limb of the above-mentioned regional anticline (Fig. 11), where it forms at least 20

stratiform lenticular orebodies. The Fe mineralisation is characterised by banded, massive,

disseminated, brecciated and veins styles, with seven recognised ore types that include:

diopside-magnetite, garnet-magnetite, diopside-amphibole magnetite,

quartz-albite-magnetite-hematite, apatite-magnetite and quartz-pyrite-magnetite. The main ore

minerals are magnetite, pyrite, chalcopyrite and pyrrhotite. The wallrocks exhibit skarn assemblages,

such as garnet, diopside, actinolite, tremolite, scapolite, epidote and chlorite. At least four stages of

skarn have been recognised (Xu et al., 2010), namely: 1) prograde stage with

clinopyroxene-garnetealbite-scapolite-apatite; 2) retrograde stage with

magnetite-clinopyroxene-garnet-amphibole-scapolite-apatite-epidote-chlorite-quartz; 3) sulphide

stage with pyrite-chalcopyrite-pyrrhotite-garnet-chlorite-quartz-calcite; and 4) supergene stage with

hematite-goethite-malachite-quartz-calcite. These four paragenetic stages conform to other Fe skarn

deposits (Pirajno, 2009). The Mengku iron skarn, and probably other skarns in the same metallogenic

belt (Fig. 4), were formed in a continental margin setting, during Early-Middle Palaeozoic subduction

under the Altay microcontinent (Yang et al., 2010). The intrusion of the Mengku granite (400 Ma),

north of orebody No. 1 (404 Ma) and the Qiongkuer granite (399 Ma) in the Mengku area into the

Kangbutiebao Formation (Yang et al., 2010), resulted in the development of skarns near and along the

contacts of the plutons, apophysis and dykes with the Kangbutiebao Formation volcanic rocks and

limestone. Following the development of these skarns, iron oxides (mostly magnetite) precipitated

from the hydrothermal fluids to form the Mengku skarn-type iron deposit.

3.3.2 Dahongshan iron-copper ore deposits

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The Dahongshan deposit is located 300 km from the city of Kunming, Yunnan Province, and is

estimated to contain ca. 350 Mt of ores with an average Fe grade of 60 wt. % after beneficiation (Qian

and Shen, 1990). The Dahongshan Group hosts the mineralization and it consists of the

Paleoproterozoic metamorphic submarine volcanic rocks and sedimentary rocks (Fig.12; Qian and

Shen, 1990).

The Dahongshan Group consists of volcanic and sedimentary rocks that were metamorphosed to

between upper greenschist and lower amphibolite facies. Metamorphic grade and intensity of

deformation varies regionally, but most outcrops show strong schistosity and some rocks are tightly

folded. The Dahongshan Group metasedimentary rocks include coarse to fine-grained siliciclastic rocks,

carbonate and volcaniclastic rocks. Siliciclastic rocks include quartzite, mica schists and polymictic

meta-conglomerates. Unimodal cross-bedding is clearly visible in quartzite, suggesting fluvial sediment

transport from a present day north-westerly direction. Volcaniclastic rocks such as volcanic breccia,

conglomerate, tuff and volcanic sandstone are found within the Manganghe and Hongshan Formations

(Greentree and Li, 2008). All carbonate units were metamorphosed to marble, with compositions

varying from pure dolomitic marble to those containing garnet or amphibole. The presence of

hornblende and garnet suggest that the protoliths contained some detrital materials. Petrogenic studies

of the metavolcanic rocks have used both major element (e.g., Qian and Shen, 1990; Hu et al., 1991)

and trace element (Xu, 1999) geochemistry. Major element geochemistry (e.g., SiO2, K2O and Na2O) is

known to be an unreliable indicator of lithology and tectonic setting in areas with complex

hydrothermal alteration and metamorphism (e.g., Pearce and Cann, 1971, 1973; Winchester and Floyd,

1977). Xu (1999) argued that the more immobile trace elements (e.g., HFSE and LREE) still preserve

the original composition of the metavolcanic rocks in the Dahongshan Group. Using the more

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immobile trace element geochemistry, Xu (1999) suggested that the volcanic rocks had a tholeiitic

composition, similar to modern mid-ocean ridge basalts (Fig. 7d).

The iron oxide and iron-copper orebodies occur in the vicinity of the volcanic center, and host at

least 43 individual iron oxide and copper (gold) mineral occurrences. Most high grade orebodies

occur in the Na-rich metamorphic volcanic rocks and in the transitional belt from volcanic rocks to

sedimentary rocks. Furthermore, petrochemical investigations have shown that the iron-copper

deposit is closely related to Na-rich volcanic rocks (Qian and Shen, 1990). Additionally, siderite

deposits have also been recognized in the metamorphic Na-rich volcanic rocks. 40

Ar/39

Ar dating of

rocks from these deposits suggests that the mineralization occurred during ca. 780-800 Ma, during a

period of plume-related magmatism on the South China Block (Greentree et al., 2006).

3.3.3 Kaladawan iron ore district

A number of iron ore deposits have been discovered in the Cambrian Zhuabulake and Simierbulake

Formations (Fig.5), such as the Baijianshan, 88, and 7918 iron deposits (Fig.13). These deposits define

an ore belt which extends 12 kilometers. Basalt is the major host rock of these deposits, interbedded

with sedimentary rocks, and exhibits aphanitic texture and massive, layered, amygdaloidal and pillow

structure. The volcanic rocks exhibit a tholeiitic differentiation trend in the SiO2-TFeO/MgO diagram

(Fig. 7c). All the deposits in this district share many similarities in their geological characteristics, such

as their conformable occurrence within basalt and marble, although some granitic intrusions are also

exposed in some deposits. Wall rock alteration assemblage predominantly includes garnet+epidote. In

the 7918 deposit for example, the iron ores occur as stratified or stratoid beds, and are in conformable

contact with the wall rocks including basalt and marble. The length of the main orebody (Fe at 41 wt.%)

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is ~580m with an average thickness of 12.5m. The ore types include banded, massive and disseminated,

and ore minerals mainly composed of magnetite.

4. Discussion

4.1. Origin of skarn in the Chinese SVIO deposits

As described in the section 3, skarn minerals, e.g. diopside and garnet, are extensively developed in

some of the Chinese SVIO deposits, and are very closely associated with iron oxide minerals. The

common feature of these deposits is that the host rocks contain carbonate. This feature has led to a

debate on the genesis of some of these deposits, especially those with considerable amount of skarn

minerals (e.g. Yang et al., 2010).

As has been recognized in earlier studies (e.g. Knopf, 1918), the formation of a skarn deposit is a

dynamic process. In most large skarn deposits there is a transition from early/distal metamorphism

resulting in hornfels, reaction skarn, and skarnoid, to later/proximal metasomatism resulting in

relatively coarse-grained ore-bearing skarn. Due to the strong temperature gradients and large fluid

circulation cells caused by magma intrusion (Norton, 1982; Salemink and Schuiling, 1987), the

formation of skarn can be considerably more complex than the simple model of isochemical

recrystallization typically invoked for regional metamorphism. For example, early metamorphism and

continued metasomatism at relatively high temperature (Wallmach et al., 1989, describe

temperatures>1200°C) are followed by retrograde alteration as temperatures decline. The shallowest

(and youngest) known skarns are presently forming in active geothermal systems (McDowell and

Elders, 1980; Cavarretta et al., 1982) and hot spring vents on the seafloor (Zierenberg and Shanks,

1983). These skarns represent the distal expression of magmatic activity, and locally, those skarns

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have some features of igneous rocks, and have been interpreted to be of magmatic origin in some

Chinese literature (e.g., Wu et al., 1996). However, the link between space and time is a common

theme in these iron ore deposits and requires careful interpretation of features which may appear to

occur only in particular localities (e.g. Barton et al., 1991).

For example, the Yamansu iron deposit contains considerable amount of stratiform skarn. Some

authors suggest that the skarns could be genetically related to a buried intrusion (e.g. Mao et al.,

2005), whereas others consider them to be related to coeval submarine volcanism (Jiang, 1983).

Recently, we conducted the laser ablation inductively coupled plasmamass spectrometry (LAICP-MS)

U-Pb zircon dating of the basalts and skarns and yielded almost coeval ages of 324.4 ± 0.94 and

323.47 ± 0.95 Ma, respectively (Hou et al., 2013). This suggests that the hydrothermal fluids that

generated the skarns could be a mixture of evolved magma-derived fluids and convecting sea water

driven by the heat from the shallow active magma chamber.

4.2. Metallogenesis of Chinese SVIO deposits

Like its subaerial counterpart such as the Kiruna style iron deposit, the origin of Chinese SVIO

deposits is uncertain and remains controversial (e.g., Jiang, 1983; Jiang and Wang, 2005). Except for

those of skarn-related origin as mentioned above, the Chinese SVIO deposits have been interpreted

variously including magmatic origin (liquid immiscibility) (e.g. Zhang et al., 1987),

exhalative-synsedimentary (Yuan, 2003), or epigenetic-hydrothermal associated with igneous

intrusion (e.g., Yang et al. 2007) or active deep-seated magma chamber (Hou et al., 2013).

Some of the SVIO deposits are clearly of magmatic origin or formed initially through magmatism,

as evidenced from their geological and geochemical features, such as high proportion of apatite in the

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magnetite ores (e.g., Abagong in Altay). Although this deposit display signature of magmatic origin

(Chai et al., this issue), but the mechanism by which the ore-bearing melt formed is still unclear

(Pirajno et al., 2009). Moreover, the model involving ore-bearing extrusive activities is mainly

inferred from the studies of the SVIO deposits occurring near the volcanic center, such as the

Chagangnuoer and Zhibo iron deposits in Western Tianshan (Wang and Jiang, 2011). Hence, a

magmatic model alone cannot entirely account for all the geological-geochemical signatures

recognized in these deposits. Specifically, many of the Chinese SVIO deposits show signatures of

hydrothermal activities, as reflected by the low temperature mineral assemblage, mainly involving

enrichment of iron in the existing iron ores, and significant presence of submarine tuff or tuffaceous

rocks in the close proximity of the orebodies. Since it has been widely accepted that the hydrothermal

minerals formed later than the magmatic ones (e.g. Hedenquist and Lowenstern, 1994), we infer that

the SVIO deposits which occur near the volcanic center probably formed initially as a result of

multi-stage and multiple processes, such as ore-bearing magma eruption, sedimentation of volcanic

pyroclastic rocks, and even exhalation-sedimentation. Most of the SVIO deposits discussed here

occurring away from the volcanic center show significant features of exhalation-sedimentation

instead of the involvement of ore-bearing magma eruption. For example, the presence of sulfide, chert

and jasper in the ores belong to Si-Fe-Mn formation which is commonly regarded as the evidence for

a seafloor-exhalation-sedimentation origin (Slack et al., 2009).

Thus, even though all the iron oxide deposits described in this paper are classified as SVIO

deposits, they probably formed by different processes related to submarine volcanism and subsequent

hydrothermal events. Therefore, the origin of the SVIO deposits occurring near the volcanic center or

vent is dominated by magmatic-hydrothermal process, whereas for those away from the center, the

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main mechanism was controlled by exhalation-sedimentation, probably aided by sea water. In

addition, most of these SVIO deposits were subsequently overprinted by metamorphism, deformation,

post-magmatic hydrothermal event and supergenesis (Jiang and Wang, 2005). For example, the

Mengku iron deposit in the Altay was influenced by post-magmatic hydrothermal event possibly

caused by the emplacement of two major intrusions of biotite granite and tonalite (Yang et al., 2010).

4.3 Relationship between nature of magmas and enrichment of iron

In general, the formation of ore-bearing magma can be attributed to magmatic differentiation,

either following a Fenner trend of differentiation (Fenner, 1929) or immiscility of iron oxide melt

(Veksler et al., 2006). As many authors have pointed out, the magma differentiation trends (Bowen or

Fenner trend) are controlled by the onset of magnetite fractionation, which is in turn is controlled by

oxygen fugacity (Osborn, 1959). Increasing oxygen fugacity (fo2) can cause marked Si enrichment

and Fe depletion in residual liquids in response to the fractionation of magnetite in the early stage

(Toplis and Carroll, 1995). In contrast, low fo2 delays the onset of magnetite crystallization leading to

prolonged Fe-enrichment in magma, exhibiting a Fenner trend (Jang, 2001). Such a trend is evident in

most of the SiO2-TFeO/MgO plots of these rocks, with tholeiitic affinity for the less evolved magmas

(Fig.7), and clinopyroxene + plagioclase fractionation is widely recognized for the basic and

intermediate volcanic rocks. These iron-rich magmas could lead to the generation of an iron oxide

fraction (ore-bearing magma) through liquid immiscibility (Veksler et al., 2006) or produce magnetite

ores by fractional crystallization (Jang et al., 2001).

Although the iron mineralization is genetically related to the magma differentiation, specific

mineralization patterns are seen in different deposits. For example, in the Abagong deposit, the

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presence of considerable amount of apatite in the massive iron ores is similar to the Kiruna deposit

(Frietsch, 1978; Nyström and Henriquez, 1994) and the ‘porphyry iron deposit’ in the Middle-Lower

Yangtze Valley in eastern China (Hou et al., 2009). Hence, the iron ores in Abagong might have

formed as a result of liquid immiscibility which was probably triggered by the enrichment of

phosphorous in the extremely evolved magma system (Suk, 1998). However, due to the lack of isotope

composition of the apatite, it is not clear whether the enrichment of phosphorous is caused by crustal

contamination (Hou et al., 2010) or the fractionation of anhydrous silicate phases increasing the

phosphorous contents in the residual magma (Green and Watson, 1982; Spengler and Garcia, 1988).

In addition to the evidences for the involvement of magmatic processes, these deposits also show

robust signatures for a hydrothermal origin, such as the extensive occurrence of low-temperature

hydrothermal minerals, e.g. albite and chlorite. This type of deposits are predominantly spatially and

temporally associated with intermediate to basic submarine volcanic rocks. Because basic and

intermediate rocks contain much higher Fe contents than felsic rocks, they could provide sufficient

iron sources for the iron mineralization. Hence, we consider large hydrothermal circulation systems,

particularly in the vicinity of volcanic center where seawater infiltrates down through fractures and

returns at high temperatures, possibly driven by the active magma chamber, could form iron-rich

fluids by leaching the relatively iron-rich volcanic rocks. For example, the extensive albite alteration

in Dahongshan deposit probably resulted from sodium alteration and Fe loss of basalts by leaching of

hydrothermal fluids (seawater-dominant) (Qian and Shen, 1990).

4.4. Significance of volcanic sedimentation

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The major involvement of subaqueous sedimentation of iron-rich material during the

mineralization could be one of the most distinguishing signatures from its terrestrial counterpart

(Carey and Sigurdsson, 2007). The submarine eruption of ore-bearing magma or iron-rich magma and

hydrothermal vent gas/fluids give rise to eruption columns that are a dispersion of gas and solid

particles containing ore-bearing brine, pumice, volcanic debris and pyroclastic. A common

assumption about submarine volcanic eruption is that the pressure of the overlying water column is

sufficient to suppress juvenile gas exsolution so that magmatic disruption and pyroclastic activities do

not occur, except at sufficiently shallow depths (e.g. Batiza and White, 2000). This depth is generally

recognized to be about 200-1000 m and less, depending on magma composition and volatile content

(c.f. Head and Wilson, 2003) and is referred to as volatile fragmentation depth (Fisher and Schmincke,

1984). Most pyroclasts will begin to fall in the immediate vicinity of the vent (within a few meters

radius) due to the negative buoyancy (Head and Wilson, 2003). Hence, SVIO deposits probably form

by the fallout of these ore-bearing or iron-rich materials to the sea floor downcurrent from the

umbrella region of submarine eruption columns (Cashman and Fiske, 1991), or fractional

precipitation of iron which had been introduced locally into the bottom water by hydrothermal

solutions of volcanic origin, and by leaching from the relatively iron-rich volcanic rocks, such as

deep-sea basaltic lavas (Bonatti and Joensuu, 1966).

However, if the ore-bearing magma is insufficiently differentiated and lithologically monotonous

with lower alkalis content, it is considered to be unfavorable to form iron-rich magma, as evident by

the absence of large-scale and high-grade iron ore deposits near volcanic center (Jiang and Wang,

2005). Instead, in this case, high-grade iron ore deposits are always recognized to be associated with

pyroclastic-sediments away from the volcanic center, such as Songhu and Shikebutai deposits. In fact,

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at the high temperature stage, regardless of its composition, the magma contains many volatile

components (Pearce and Peate, 1995). With effusive activity, lavas rich in volatiles such as F, Cl and

CO2, are discharged. Volcaniclastic formation depends on several factors, including magma

composition, volatile concentration, eruption depth and rate and magma-water interaction

mechanisms (e.g. Gamberi, 2001; Orton 1996; Carey 2000; Fisher and Schmincke, 1994). The

processes that might be responsible for this fragmentation are magmatic explosivity, contact-surface

steam explosivity, bulk interaction steam explosivity, cooling-contraction granulation, or any

combination of these (c.f. Fouquet et al., 1998). Nevertheless, these activities lead to the breakdown

of primary volcanic rocks.

It has been found that suspended matter is typically enriched in Fe (e.g. Ferguson and Lambert,

1972). It is believed that most of the Fe, SiO2 and Mn entering sea water in hydrothermal solutions

precipitates as colloidal SiO2 and hydrated Fe (Mn) oxides which are advected by bottom currents and

deposited as crusts and sediments (Toth, 1980). Therefore, the deposition of Fe in ore concentrations

could occur at considerable distances from the volcanic vent or center (Lisitzin, 1996). Since the

concentration occurs in calm and depressed areas, it can be inferred that euxinic to oxidizing basin

environments is favorable for the formation of these deposits which are in association with clastic and

pelagic sediment, tuff, volcanic rocks and a variety of clay minerals (Mottl, 1983; Yang and Scott,

1996; de Ronde et al., 2005). However, such mechanism is still inadequate to explain the formation of

high grade iron ores because the suspended matter generally contains other particles or minerals such

as carbonates and silicates. Hence, the high grade iron ores require more efficient mechanism for the

enrichment of iron content. For example, leaching of the pyroclastics and debris relatively enriched in

iron by deep sea water or hydrothermal circulation could be a major source of iron for these high

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grade ores (e.g. Brusnitsyn and Zhukov, 2012). This process also involved the alteration of pyroclastic

rocks under the influence of hot water and steam on a large scale (Fontboté, 1990). The principal

factors determining the extent is temperature and the sanity of leaching fluids (Dekov et al., 2010).

4.5 Volcanic facies and metallogenic model

As previously studied (e.g., Chen et al., 2011; Cui et al., 2010; Chai et al., 2009; Greentree et al.,

2006), the Chinese SVIO deposits might have formed along divergent plate boundaries and in

intraplate areas, such as island arc, back-arc basin and rift etc. Under the submarine environment in

the different tectonic settings, in addition to effusive flows, submarine eruptions can produce

pyroclastic deposits (e.g. composed of ‘solid fragments ejected from volcanoes’; Cashman et al., 2000)

and hyaloclastic deposits (e.g. consisting of ‘fragments of volcanic glass formed by non-explosive

shattering’; Batiza and White, 2000). During eruptions, large volumes of lava, pyroclastic and

hydroclastic sediment are released far more rapidly than any process of production of epiclastic

particles (Houghton and Landis, 1989). The episodic nature of eruptions may profoundly disrupt flow

and sedimentary environments and processes resulting in rapid changes in the depositional systems

through time. Therefore, the host rocks comprise compositionally and texturally diverse lavas and

pyroclastic rocks, most of which were emplaced in submarine environments and distribute around the

volcano core or volcanic center at different distances (Williams and McBirney, 1979). Thus, the

volcanic facies architecture reflects the contrasting character and geometry of primary volcanic and

pyroclastic facies which are strongly controlled by eruption style and emplacement processes (Fisher

and Schmincke, 1984) and the related mineralization.

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As shown in Fig. 14, the volcanic facies vary according to their closeness to the source, i.e. central,

proximal and distal facies (Williams and McBirney, 1979). For example, the rocks of central facies

are always recognized by lava domes and thick, banded lavas, lag-fall breccias of pyroclastic flows,

abundant dykes and sills, circular to elongate stocks, breccia pipes and hydrothermally altered rocks.

The shapes of orebodies are controlled by the submarine volcanic edifice (Fig.14). The proximal

facies rocks around a volcanic center deposited from pyroclastic flows, lava flows, debris

flows/avalanches, fallout processes and their erosional products. As distance from the source

increases within this facies, there is an increase in the amount of re-sedimented epiclastic and

pyroclastic debris. Particularly, the pyroclastic flow units (main body) in this zone are commonly

underlain by surge deposits (lens-like) and overlain by fine-bedded ash deposits, and block and ash

flows from dome collapse formed monolithic, massive, poorly sorted clastic debris which is reworked

by seawater, and contain debris avalanche deposits-mounds (block facies) and more normal laharic

material. The distal zone is the base of volcano and beyond. Therefore, rocks here are characterized

by a much greater lateral continuity than those of the proximal and central facies. Finely bedded

tephra composed dominantly of fine-coarse ash, outward increasing ratio of glass to crystals are

recognized in this zone where the pyroclastic flows will be thinner here than in proximal areas, and no

surge deposits, ash fall commonly occur above flows.

Moreover, considering the involvement of sea water during the mineralization process, with

increasing distance from the center of the volcano, the dominant mechanism by which iron oxide

enrichment occurs in the deposit changed from magmatic, hydrothermal to sedimentary. Consequently,

the characteristics of iron ore deposit have also perceptibly changed. For example, the iron ore deposit

occurring within lava has been mainly discovered in the central zone (e.g. Abagong and Yamansu),

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and seldom recognized in the distal zone (Jiang, 1983). The main reasons for this are: the massive

orebodies were formed by immiscible oxide melt separated from the silicate melt within crust-level

magma chamber beneath volcano center, and the brecciated ores could be attributed to the eruption of

ore magma and an explosion of the magmas at the volcano center near seafloor responsible for fluid

exsolution developed by decompression and rapid condensation. The volcanic central zone is defined

as the area overlain by lava and coarse-grained pyroclastic rocks rather than the location of the

volcanic vent (Williams and McBirney, 1979). The active magma chamber, which occurs as

subvolcanic edifice presently, is difficult to identify as the original feature had been more or less

changed by the subsequent tectonic activity. Presently, the central facies rocks could serve as a

potential surface indication of potentially economic SVIO mineralization. For example, the

deep-seated subvolcanic rocks are possibly host rocks for contact metasomatism (i.e. Fe-skarn)

mineralization, especially where the intrusions were emplaced into the carbonate strata (e.g. Einaudi,

1981). In contrast, the proximal and distal zones are dominated by fine-grained pyroclastics and

volcanic sediments. The iron ores in these zones mainly are hosted in volcanic sediments and volcanic

sedimentary-volcanic hydrothermal (Fig. 14), where the banded ores hosted in the well sorted

volcanic sedimentary rocks, such as tuffaceous rocks are commonly seen. However, the origin of

these deposits is complex due to the episodic nature of eruptions (c.f. Hildenbrand et al., 2008) which

could also lead to the development of ephemeral subvolcanic magma reservoirs (Zellmer et al., 2005).

Thus, these volcanic processes, combined with the related hydrothermal activities and transformation

led to a complex metallogenesis for the SVIO deposits. For the SVIO deposits located near plate

boundary, subsequent tectonic processes such as regional metamorphism (medium to high grade), and

contact metamorphism probably played important roles. In the proximal and distal zones, the

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involvement of sea water plays a more important role associated with the deformation and

metamorphism, resulting in changes in the shape, nature of metal distribution and types. Therefore,

our genetic model which correlates the origin of these deposits with the space-time evolution of the

submarine volcanoes and principal volcanic facies variation offers a better understanding of the

metallogenesis of SVIO deposits, aiding in their further exploration in China and around the world.

5. Conclusions

The submarine volcanic iron oxide deposits are one of the most important base–metal ore deposits

in China, and typically occur within or near the paleo-seafloor in submarine volcanic environments.

These deposits are hosted in subvolcanic intrusion, lava, volcanic pyroclastic and

volcaniclastic-sedimentary rocks, or a combination of these. The iron orebodies hosted in different

volcanic facies exhibit different signatures and reflect their closeness to the volcanic center. Thus, the

iron ores formed by ore magma eruption are predominantly discovered in the vicinity of volcanic

center. Most of these deposits are characterized by widely developed skarns, which could be

interpreted as a distal expression of magmatic activity and exposed igneous rocks. Metamorphism and

continuous alteration at relatively high temperature was followed by retrograde alteration as

temperatures declined.

Geological and geochemical evidence suggest that these deposits were formed as a result of

continuous submarine magmatic activities including the subaqueous volcanic explosions, lava

eruption, volcano-sedimentary processes, and related post-magmatic hydrothermal activities. In

combination with their geological characteristics, geodynamic mechanisms and metallogenesis, we

propose a genetic model in which the origin of these deposits is related to the space-time evolution of

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the submarine volcanoes. We integrate the deposits by principal volcanic facies variation into central,

proximal and distal facies providing further insights into their metallogenic history and exploration

potential.

Acknowledgements

Financial support for this work was supported by Projects 2012CB416806 of the State Key

Fundamental Program (973), Special Fund for Scientific Research in the Public Interest

(200911007-25), Fundamental Research Funds for the Central Universities National Natural Science

Foundation of China (No.40925006), and 111 Project (B07011).

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Figure captions

Fig.1. Distribution of Chinese submarine volcanogenic iron oxide deposits (base map modified

from Zhao et al., 2004).

Fig.2. Geological map of the Tianshan Orogen showing the Awulale Metallogenetic Belt

(modified after Zhang et al., 2012), showing the locality of the several submarine volcanic iron ore

deposit.

Fig.3. The tectonic framework and distribution of iron ore deposits in the Eastern Tianshan

Mountains (modified from Wang et al., 2006).

Fig.4. Regional geological map of the Altay orogeny in NW China (Xinjiang) and distribution of

the iron, gold and base metal deposits ores (modified from Xu et al., 2010).

Fig.5. Geological map and distribution of iron deposits in the Kaladawan area, eastern part of

the Altyn Tagh Mountain (modified from Chen et al., 2009).

Fig.6. Geological map showing the Precambrian rocks and distribution of iron ore and copper

deposits in the southwestern margin of Yangtze Craton (modified from Qian and Shen, 1993).

Fig.7. SiO2-TFeO/MgO diagrams showing the magmatic differentiation of iron ore hosting rocks.

Data source: Chagangnuoer (Wang and Jiang, 2011), Altay (Mengku+Abagong; Zhang et al., 1987),

Kaladawan (Cui et al., 2010), Dahongshan (Qian and Shen, 1990).

Fig.8. Geological map (a) and cross section of Yamansu Fe-Cu deposit (modified from Mao et

al., 2005).

Fig.9. Geological map (a) and cross-section (b) of Abagong iron (apatite) deposit in Altay,

Xinjiang (after Yang et al., 2011).

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Fig.10. Geological map and the insert maps are plane (a) and cross section of No.1 orebody in

Songhu iron deposit in Awulale Metallogenic Belt in Western Tianshan Mountains (modified from

Shan et al., 2009).

Fig.11. Geological map (a) and cross section (b) of Mengku iron deposit in Altay, Xinjiang

(modified from Xu et al., 2010).

Fig.12. Geological map of Dahongshan Fe-Cu deposit, Southwestern margin of Yangtze Craton

(modified from Qian and Shen, 1990). The numbers with circle indicate the numbers of the faults.

Fig.13. Geological map (a) and cross section (b) of 7918 iron deposit in Kaladawan area, eastern

part of the Altyn Tagh Mountain (modified from Chen et al., 2009).

Fig.14. Proposed genetic model for the iron ore deposits based on facies variations in submarine

volcanic rocks from a large central vent composite volcano (modified from Williams and McBirney,

1979). Central zone is also known as the vent facies. Products of each zone/facies are listed in the

illustration.

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