Oceanic Hotspots

Intraplate volcanism occurs in both submarine and subaerial regions on the Earth’s surface. However, most of the magmatic activity responsible for the construction of the Earth’s volcanic structures takes place under the sea in oceanic basins and along spreading centers. The divergent plate boundaries (Mid-Ocean Ridges) and back-arcs are relatively straight and structurally continuous features, making them easier to investigate than the more scattered and dispersed intraplate volcanic provinces within the oceanic basins.

The islands of French Polynesia were discovered and populated by Polynesians between 500 B.G. and A.D. 500. European exploration of the region began in the seventeenth century. The main island groups (Society, Marquesas, Tuamotu, and Austral) have long been known, but the submarine topography has only been explored in the past fifty years using conventional echo sounders, and only in the past twenty years using higher-resolution multibeam sonar systems.

Since the oceanic column provides an optical, thermal, and to a large extent, chemical shield for the remote sensing of the planet’s surface, careful monitoring of seismic activity on the ocean floor remains one of the few methods of studying submarine volcanic activity. This line of research goes back to more than fifty years ago, when based on a suggestion of (1946), (1954) obtained a detailed history of the 1952 eruption of the Myojin Volcano, south of Japan, using teleseismic T waves propagated in the oceanic column over a distance of 8600 km to a receiving array on the California coastline.

The concept of “continental drift” advocated by A. Wegener from about 1912 until his death in 1930 was based on geological observations and (for those times) modern principles of isostasy. His idea about the “wandering continents” was complemented and strengthened later on by important notions, among which were the rejuvenation of the oceanic lithosphere and its absorption in the inner Earth after subduction that later evolved as the theory of “plate tectonics” and gained general acceptance in the 1970S. Plate tectonics introduced the radically new notion in geodynamics of the large horizontal motion of about 100 km thick lithospheric plates that were gliding on their substratum, the asthenosphere. The dynamics of the plates, thought to be driven by convection currents in the asthenosphere, determines the relief of the Earth’s crust.

The Hawaiian-Emperor Seamount chain (H-E SMC) on the Pacific Plate Figs. 4.1 and 4.2 a,b) is the best-defined hotspot track on the Earth. If hotspots are surface manifestations of deep, fixed sources of mantle plumes (Morgan 1971, 1981), then the along-track volcanic age progression away from Hawaii (e.g., Clague and Dalrymple 1989) must record the direction, absolute velocity, and possible changes of the Pacific Plate motion. This would suggest that the prominent ∼43 Ma Bend along the H-E SMC reflects a sudden change in Pacific Plate motion direction by ∼60°. However, the actual cause of the 43Ma Bend is unknown. A leading hypothesis is that the collision between India and Eurasia some ∼45 Ma ago might have triggered the sudden reorientation of the Pacific Plate motion from northward to northwestward, hence the 43Ma Bend (Dalrymple and Clauge 1976; Patriat and Achache 1984).This collision, however, is shown to have had no effect on the Pacific Plate motion (Lithgow-Bertelloni and Richards 1998). The lack of apparent mechanism for such a sudden change in Pacific Plate motion direction led to the speculation (Norton 1995) that the ∼43 Ma Bend may have resulted from a southward drift of the Hawaiian hotspot prior to ∼43 Ma. Indeed, recent paleomagnetic studies (Tarduno and Gee 1995; Tarduno and Cottrel 1997; Christensen 1998; Sager 2002), plate reconstructions (Acton and Gordon 1994; Norton 1995, 2000; DiVenere and Kent 1999; Raymond et al. 2000), mantle flow models (Steinberger and O’Connell 2000), and statistical analysis of plate motions using seamount geochronology (Koppers et al. 2001) tall indicate that hotspots are not fixed, but they move individually or in groups at speeds up to 60 mm yr⁻¹.

Intraplate volcanic activity in the Pacific (Fig. 5.1) has been observed and/or seismically recorded in the following areas: (1) Samoa Islands with the Rockne Seamount (Johnson 1984); (2) the Society archipelago with the Teahitia Seamount (Cheminée et al. 1989; see Sect. 2.3.5) and seismically active Mehetia Island (Talandier and Okal 1987; see Sect. 2.3.5.1); (3) the Austral archipelago with Macdonald Seamount (Johnson 1970; see Sect.2.4.1); (4) the Pitcairn hotspot with the Bounty Volcano (Stoffers et al.1990); and (5) Hawaii. Numerous geological and geophysical investigations have been carried out in the Pacific, including the Line Islands (Schlanger et al. 1984), Marquesas Islands (McNutt 1989; Duncan and McDougall 1974), Austral Society Islands (Duncan and McDougall 1976) and the Louisville chain (Lonsdale 1988; Watts et al.1988). Nevertheless, except for Hawaii (Macdonald and Abbott 1970; Fornari et al. 1979, 1980, 1988; Lonsdale 1989), little is known about the morphology and structure of submarine intraplate volcanoes in the Pacific.

Landslides are common features of oceanic islands and playa key role in their evolution. Caused by caldera collapse or flank collapses, they can be classified into three types: (1) rock falls, (2) slumps or (3) debris avalanches (Moore et al. 1989). Rock falls, or superficial landslides, are mainly related to erosion processes of the subaerial parts of the island. The pieces of debris are less than 1 m in size, and their surface is rippled. Flank collapses generally produce giant submarine landslides, with a horseshoe-shaped feature at their head (Moore et al. 1989). The landslides due to a deep listric fault are cataclysmic events producing fast moving debris avalanches. Deposits can extend over several hundred kilometers away from an island and are characterized by thicknesses less than 2 km, with a hummocky terrain at their lower part. Side-slip over deep fault is termed slump Fig. 6.1. Slumps are slow-moving slope instabilities. The thickness of the deposits can be as much as 10 km, since the primitive volcano flank is less shattered and disrupted than in the case of a debris avalanche. The causes of major lateral collapses are still a matter of debate, but in most cases they are thought to be related to magma intrusion in the rift zones (Denlinger and Okubo 1995; Keating and McGuire 2000).

Basaltic volcanism mainly occurs in three tectonic settings on the Earth. Volcanism along sea-floor spreading centers produces Mid-Ocean Ridge basalts(MORB) that are depleted in incompatible elements. Volcanism above intra-oceanic subduction zones produces island arc basalts (IAB) that are enriched in water-soluble incompatible elements (e.g., Ba, Rb, Cs, Th, U, K, Pb, Sr), but depleted in water-insoluble incompatible elements (e.g., Nb, Ta, Zr, Hf, Ti). MORB and IAB are products of plate tectonics, and their geochemical differences result from differences in their respective sources and physical mechanisms through which they form. MORB are formed by plate-separation-induced passive mantle upwelling and decompression melting, thus sampling the uppermost mantle that is depleted in incompatible elements. Depletion of the MORB mantle is widely accepted as resulting from the extraction of incompatible element-enriched continental crust during the Earth’s early history (Armstrong 1968; Gast 1968; O’Nions and Hamilton 1979; Jacobsen and Wasserburg 1979; DePaolo 1980; Allègre et al. 1983; Hofmann 1998). IAB are widely accepted as resulting from subducting slab-dehydration-induced melting of mantle wedge peridotites, giving rise to the characteristic geochemical signatures of slab “component”, which is rich in water and water-soluble elements (e.g., Gill 1981; Tatsumi et al. 1986; McCulloch and Gamble 1991; Stolper and Newman 1994; Hawkins 1995; Pearce and Peate 1995; Davidson 1996).

The South Pacific is characterized by a large number of active hotspots, many of which have been active for long periods of time (possibly as long as 120 Ma, Staudigel et al. 1991) producing extensive island and/or seamount chains (see Introduction, Fig. 0.1. The hotspots are presently located either beneath relatively old lithosphere (e.g., Society, Pitcairn, Australs, Marquesas, Juan Fernandez) or lie closer to the spreading axis (Foundation, Easter) (see Sect. 5.1). During the last fifteen years, the German-French initiative to study these hotspots has resulted in a vast amount of petrological and geochemical data being collected on fresh, mainly submarine volcanics.

Geological processes are consequences of the Earth’s thermal evolution. Plate tectonics, which explain geological phenomena along plate boundaries, elegantly illustrate this concept. For example, the origin of oceanic plates at ocean ridges, the movement and growth of these plates, and their ultimate consumption back into the Earth’s interior through subduction zones provide an efficient mechanism to cool the Earth’s mantle, leading to large-scale mantle convection. Mantle plumes, which explain another set of global geological phenomena, cool the Earth’s deep interior (probably the Earth’s core) and represent another mode of Earth’s thermal convection (e.g., Davies and Richards 1992). Plate tectonics and plume tectonics are thus genetically independent from each other, However, when the rising plumes approach the lithospheric plates, interactions between the two inevitably result. Such interactions are most prominent near ocean ridges, where the litho-sphere is thin and the effect of mantle plumes is best revealed. “Plume-ridge inter-action” has been a hot topic in recent years, and much effort has been expended in this area aimed at understanding the geological, geochemical, and geodynamic consequences (Schilling et al. 1983, 1994, 1995, 1996, 1999; Schilling 1991; Feighner and Richards 1995; Ho and Lin 1995a,b; Ito et al. 1996; Kincaid et al. 1995, 1996; Ribe 1996; Sleep 1996; Haase and Devey 1996; Hekinian et al. 1996, 1997, 1999; Pan and Batiza 1998; Niu et al. 1999; Graham et al. 1999; Maia et al. 2000; Georgen et al. 2001; Haase 2002).

The Macdonald Seamount Fig. 10.1a is located at the tip of the Austral hotline (Johnson 1970, 1980; Talandier and Okal 1984). The activity of the Austral and Society hotspots has been closely monitored by the detection of seismic swarms recorded by the French Polynesian seismic network (“Réseau Sismique Polynésien”, RSP). Since the Austral Islands are too far away from the receiving stations, only ‘T’ waves have been detected from the Macdonald Seamount (Talandier and Okal 1984; see Sects. 2.4.1 and 2.4.2). This seamount is one of the most active submarine volcanoes in the world (Cheminée et al. 1991) and was first noticed after a strong seismic swarm was detected by the hydrophones of the Hawaiian Institute of Geophysics Network (Norris and Johnson 1969). A multibeam bathymetric survey (NO JEAN CHARCOT, FS SONNE and NO L’ATALANTE) of the most recent seamounts of the Society and the Austral hotspots was undertaken in 1986 and 1987. The edifices were also sampled, and several dredge hauls were undertaken on top of the Macdonald Seamount Fig. 10.1b. Related publications have mainly dealt with the morphology and the structure of the Society and Austral hotspots and the petrology of the volcanics (Stoffers et al. 1989; Sect. 5.3.1.1; Hekinian et al. 1991). Among the samples recovered from the Macdonald Seamount, highly vesicular pillow lavas, volcaniclastics and accidental rock debris were found. The gabbroic clasts were ejected during hydromagmatic explosive events nearly twenty years after the seamount was first discovered. Later, they were partially covered by basanite lapilli during further explosions (Sect. 5.3.1.1; Hekinian et al. 1991).

The Foundation Chain was first detected using a combination of satellite altimetric and conventional geophysical data (Sandwell 1984, Mammerickx 1992) and described initially as a ∼1350 km long chain of seamounts trending approximately in the direction of the absolute motion of the Pacific Plate (Mammerickx 1992)Fig. 11.1. A significant section of the Foundation Chain lies in a tectonic setting influenced by a change in the direction of sea-floor spreading between 26 and 11 Ma (Herron 1972; Lonsdale 1988; Mayes et al. 1990; Mammerickx 1992). This motion change is reflected in the curvature of the Agassiz Fracture Zone (FZ) and its west-east shift in orientation between the Resolution/Del Cano and Chile FZs Fig. 11.2. A segment of the Nazca plate was transferred to the Pacific Plate during this period of reorganization to form the short-lived Selkirk Microplate Fig. 11.2 (Mammerickx 1992; Tebbens and Cande 1997; Tebbens et al. 1997) via a spreading-ridge propagation event between chron 6C (23.4–24 Ma) and Chron 6(0) (20.2 Ma) (Tebbens and Cande 1997).

Submarine iron and manganese deposits have a widespread occurrence in the oceanic environment. Genetically they can be subdivided into three discrete types (Boström 1983; Usui and Terashima 1997): (1) hydrogenetic, (2) diagenetic, and (3) hydrothermal. Hydrogenetic deposits occur as crusts on seamounts and other volcanic outcrops and as nodules on abyssal sediments via the direct precipitation of ironmanganese oxides and hydroxides from seawater (Koschinsky and Halbach 1995). Since these oxides and hydroxides have high adsorption capabilities, hydrogenetic crusts are characterized by relatively high trace element contents (e.g., Pb, Co, Ni) and slow growth rates (on the order of mm Ma⁻¹; Segl et al. 1989). Mineralogically, they are composed of vernadite (Fe-rich δ-MnO₂) and X-ray amorphous iron oxyhydroxides (δ-FeOOH) (Hein et al. 1999). The growth and the composition of diagenetic iron-manganese nodules are controlled by diagenetic element supply from the sediments. These nodules are characterized by high growth rates (on the order of 10–200 mm ka⁻¹) and high Mn/Fe ratios as well as low trace element contents (Reyss 1982).The third type of deposit, the hydrothermal iron-manganese crust, is ubiquitous along Mid-Oceanic Ridges and back arc spreading centers. They are characterized by high growth rates (cm ka⁻¹) and low trace element content. Their origin is closely related to the emanation of metal-rich hydrothermal fluids. These fluids are the result of hydrothermal convection cells that are fueled by the heat of a subsurface magma.

In the past, marine hydrothermal systems were studied by numerous work groups that focused on different research aspects. For instance, mechanisms of fluid -associated particle transport and diverse hydrothermal mineralizations have been studied by von (1987), (1992), (1993), (1993), (1994), (1995), (1997), Scholten et al. (see Sect. 12.2) and others. Moreover, biologists found that hydrothermal vent systems can host chemosynthetic organisms (Tunnicliffe 1991; Jannasch 1995; Nelson and Fischer 1995; Dando et al. 1995). Hydrocarbons observed in hydrothermal systems were ascribed to abiogenic and biogenic formation processes. In particular, numerous hydrothermal vents on deep-seated sea floor show CH₄ of abiogenic origin (Welhan 1988; Charlou et al. 1996, 2002). Hydrocarbon formation could occur via diverse processes like water-/rock reactions (including serpentinization processes) investigated by (1980), (1985), (1995), and (2002). Moreover, hydrothermal trace gases can also be introduced by mantle emanations (Craig and Lupton 1981; Welhan 1988) or formed in the Earth’s crust by thermocatalytic (Simoneit 1983; Michaelis et al. 1990) or abiogenic reactions (Apps 1985; Sherwood Lollar et al. 1993). On the other hand, biogenic methane formation (and/or hydrocarbon oxidation) may be caused by microbial activities within the vents and at or near the sediment surface at temperatures below 113 °C (Huber et al. 1990; Burggraf et al. 1990). These biogenic gases of relatively shallow origin may be superimposed on hydrothermal trace gas components formed by abiogenic reactions in the Earth’s crust.

Underwater volcanism is the predominant phenomenon taking place on Earth. About 71% of the Earth’s sea-floor surface is the site of volcanism, which mainly extrudes basaltic rocks. The mineralogical and chemical compositions of oceanic rocks are important factors that affect the formation of the oceanic crust, because the rock’s composition influences its rate of extrusion, the morphology of lava flows, the cooling rate, the mode of emplacement, etc. The magmatic history of the various geological provinces is also controlled by the processes of partial melting of a heterogeneous mantle and of crystal-liquid fractionation within the magmatic reservoir. Thus, the magmatic history (the rocks’ petrology) can be better understood if we first clearly define and understand both the mineral and chemical composition and the morphology of rocks found in different geological settings and provinces of the ocean floor.