EARTH AS AN EVOLVING PLANETARY SYSTEM Part 8

The Supercontinent 9 Cycle and Mantle-Plume Events Introduction Supercontinents have aggregated and dispersed several times during geologic history, although our geologic record of supercontinent cycles is only well documented for the last two cycles: Gondwana–Pangea and Rodinia (Hoffman, 1989; Rogers, 1996). It is generally agreed that the supercontinent cycle is closely tied to mantle processes, including both convection and mantle plumes. However, the role that mantle plumes may play in fragmenting supercontinents is still debated. Condie (1998) and Isley and Abbott (1999) have presented arguments that mantle-plume events have been important throughout the Earth’s history and may account for the episodicity of continental growth as described in Chapter 8. Although the meaning of mantle-plume event varies in the scientific literature, I shall constrain the term to refer to a short-lived mantle event (£100 My) during which many mantle plumes bombard the base of the lithosphere. During a mantle-plume event, plume activity may be concen-trated in one or more mantle upwellings, as during the mid-Cretaceous mantle-plume event some 100 Ma, when activity was focused mainly in the Pacific mantle upwelling. However, as I have pointed out (Condie, 1998; Condie, 2000), alleged Precambrian mantle-plume events at 2.7 and 1.9 Ga correlate with maxima in worldwide production rate of juvenile crust; thus, these events may not have been confined to one or two mantle upwellings. One of the first models presented to explain episodic continental growth was that of McCulloch and Bennett (1994). They proposed a nonrecycling model involving three reservoirs: continental crust, depleted mantle, and primitive mantle. It assumes that the volume of depleted mantle increases with time in a stepwise manner, which is linked to major episodes of continental crust formation at 3.6, 2.7, and 1.8 Ga. The isotopic and trace element composition of the upper mantle is buffered by the progressive extraction of continental crust and the increasing size of the depleted mantle reservoir. 315 316 The Supercontinent Cycle and Mantle-Plume Events Stein and Hofmann (1994) were among the first to advocate that episodic instability at the 660-km seismic discontinuity controls the growth of continental crust. They suggested that convection patterns changed in the mantle from layered convection (the normal case), when the growth rates of continental crust were relatively low, to whole-mantle convection when the growth rates were high. Whole-mantle convection occurs in short-lived episodes during which subducted slabs accumulated at the 660-km disconti-nuity catastrophically sink into the lower mantle in a manner similar to that proposed by Tackley et al. (1994). One of the important features of the Stein-Hofmann model is that during periods of whole-mantle convection, plumes rise from the D” layer above the core and replenish incompatible elements to the upper mantle, which has been depleted by oceanic crust and arc formation. Based on the same theme of instability at the 660-km discontinuity and using parameterized mantle convection, Davies (1995) proposed catastrophic global magmatic and tectonic events at a spacing of 1 to 2 Gy. The favored models show layered convection, which becomes unstable and breaks down episodically to whole-mantle convection as in the Stein-Hofmann model. During the catastrophic mantle overturns, hot lower mantle material is transferred to the upper mantle and may be responsible for rapid episodic growth of juvenile crust, as well as for replenishing the upper mantle with incompatible elements. Peltier et al. (1997) extended thermal constraints to more thoroughly evaluate the catastrophic mantle models. These investigators quantified the physical processes that control the Rayleigh number at the 660-km discontinuity, which in turns controls the frequency of slab avalanches at this discontinuity. They also suggested a correlation between the avalanche events and the supercontinent cycle. Their results imply that slab avalanches occur at a spacing of 400 to 600 My and that they are brought about by the growth of an instability in the thermal boundary layer at the 660-km discontinuity. During and after slab avalanches, a large mantle downwelling is produced directly above the avalanches; this downwelling attracts fragments of continental lithosphere, thus leading to the formation of a supercontinent. Based on the episodic occurrence of juvenile crust and associated mineral deposits, Barley et al. (1998) proposed a global tectonic cycle beginning in the late Archean with the breakup of a supercontinent. Enhanced magmatism from 2.8 to 2.6 Ga results from a global mantle-plume event. I also proposed (Condie, 1998) a model to explain the episodic growth of juvenile crust based on episodic mantle-plume events, which will be described in more detail later in this chapter. Supercontinent Cycle The most detailed and extensive coverage of the supercontinent cycle comes for the recent supercontinent Pangea. Pangea at 200 Ma was centered approximately over the African geoid high (Fig. 4.4a), and the other continents moved from this high during the breakup of Pangea. Because the geoid high contains many of the Earth’s hotspots and is characterized by low seismic-wave velocities in the deep mantle, it is probably hotter Supercontinent Cycle 317 than average, as explained in Chapter 4. Except for Africa, which still sits over the geoid high, continents seem to be moving toward geoid lows that are also regions with rela-tively few hotspots and high lower-mantle velocities, all of which point to cooler mantle (Anderson, 1982). These relationships suggest that supercontinents may affect the thermal state of the mantle, with the mantle beneath continents becoming hotter than normal, expanding, and producing the geoid highs (Anderson, 1982; Gurnis, 1988). This is followed by increased mantle-plume activity, which may fragment supercontinents or at least contribute to the dispersal of cratons. Supercontinent Cycle in the Last 1000 Million Years In the last 1000 My, three supercontinents have come and gone. The Meso- and Neoproterozoic supercontinent Rodinia formed as continental blocks collided primarily along what today is the Grenville orogen, which extends from Siberia along the coasts of Baltica, Laurentia, and Amazonia into Australia and Antarctica (Hoffman, 1991; Condie, 2002a) (Fig. 2.28). Gondwana formed chiefly between 600 and 500 Ma (Fig. 2.27), and Pangea formed between 450 and 300 Ma (Fig. 2.26). Rodinia Although Rodinia appears to have assembled largely between 1100 and 1000 Ma (Fig. 9.1), some collisions, such as those in the northwest Grenville orogen (eastern Canada) and collisions between the South and Western Australia plates (Rivers, 1997; Condie, 2003b; Meert and Torsvik, 2003; Pesonen et al., 2003) began as early as 1300 Ma. Relatively minor collisions between 1000 and 900 Ma, collisions such as Rockall–Amazonia and Yangtze–Cathaysia, added the finishing touches on Rodinia. Paleomagnetic data suggest that with the exception of Amazonia most or all of the cratons in Africa and South America were never part of Rodinia (Kroner and Cordani, 2003). These latter cratons, however, remained relatively close to each other from the Mesoproterozoic onward. Rodinia began to fragment from 800 to 750 Ma with the separation of Australia, east Antarctica, south China, and Siberia from Laurentia. Extensive dyke swarms emplaced at 780 Ma in western Laurentia may record the initial breakup of Rodinia in this area (Harlan et al., 2003). Although most fragmentation occurred between 900 and 700 Ma, the opening of the Iapetus Ocean began about 600 Ma with the separation of Baltica– Laurentia–Amazonia. In addition, small continental blocks, such as Avalonia–Cadomia and several blocks from western Laurentia, were rifted away as recently as 600 to 500 Ma (Condie, 2003b). As described in Chapter 8, Sr isotopes of marine carbonates, as proxies for seawater, can be useful in tracking the history supercontinents. As an example, consider Rodinia. It would appear that the increase in the Sr isotopic ratio of marine carbonates between 1030 and 900 Ma records the last stages in the formation of Rodinia (Fig. 9.2). The Sr isotopic ratio decreases in seawater from about 0.7074 at 900 Ma to a minimum of 0.706 from 850 to 775 Ma (Jacobsen and Kaufman, 1999). This dramatic decrease 318 The Supercontinent Cycle and Mantle-Plume Events Figure 9.1 Distribution of rifting and collisional ages used in the construction of supercontinent cycles in the last 1 Gy. Fm, formation; SC, supercontinent. Data references in Condie (2002a). COLLISIONS 6 4 2 1400 1200 1000 800 600 400 200 0 Rodinia Fm Gondwana Fm Pangea Fm New SC? Pannotia Rodinia Breakup Pangea Breakup RIFTING 6 4 2 1400 1200 1000 800 600 400 200 0 Age (Ma) 0.710 Growth of 0.709 Gondwana from Pan-African collisions 0.708 Rifting of Rodinia 0.707 Formation of Rodinia 0.706 0.705 500 600 700 800 900 1000 1100 1200 1300 Age (Ma) Figure 9.2 Distribution of the 87Sr/86Sr ratio in seawater from 1000 to 400 Ma. Points represent published data from the least altered marine limestones. Modified from Condie (2003b). Supercontinent Cycle 319 probably records the breakup of Rodinia with increased input of mantle Sr accompany-ing the breakup. The minimum is followed by a small but sharp increase in radiogenic Sr, leveling off between about 700 and 600 Ma. This small increase may reflect some of the early plate collisions in the Arabian–Nubian shield and elsewhere. The most signifi-cant change in the Sr isotopic ratio of Neoproterozoic seawater occurs between 600 and 500 Ma when the 87Sr/86Sr ratio rises to near 0.7095 in only 100 My. This rapid increase corresponds to the Pan-African collisions leading to the formation of Gondwana. As collisions occurred, land areas were elevated and a greater proportion of continental Sr was transported into the oceans. Gondwana and Pangea The formation of Gondwana immediately followed the breakup of Rodinia with some overlap in timing between 700 and 600 Ma (Fig. 9.1). The short-lived supercontinent Pannotia, which formed as Baltica, Laurentia, and Siberia briefly collided with Gondwana between 580 and 540 Ma (Dalziel, 1997), assembled and fragmented during the final stages of Gondwana construction. Pangea began to form about 450 Ma with the Precordillera–Rio de la Plata, Amazonia–Laurentia, and Laurentia–Baltica collisions (Li and Powell, 2001) (Fig. 9.1). It continued to grow by collisions in Asia, of which the last major collision produced the Ural orogen between Baltica and Siberia about 280 Ma. It was not until about 180 Ma that Pangea began to fragment with rifting of the Lhasa and west Burma plates from Gondwana. Major fragmentation occurred between 150 and 100 Ma, with the youngest fragmentation—that is, the rifting of Australia from Antarctica—beginning about 100 Ma. Small plates, such as Arabia (rifted at 25 Ma) and Baja California (rifted at 4 Ma) continue to be rifted from Pangea. Although often overlooked, there are numerous examples of continental plate collisions that paralleled the breakup of Pangea. Among the more important are the China–Mongolia–Asia (150 Ma), west Burma–Southeast Asia (130 Ma), Lhasa–Asia (75 Ma), India–Asia (55 Ma), and Australia–Indonesia (25 Ma) collisions. In addition, numerous small plates collided with the Pacific margins of Asia and North and South America between 150 and 80 Ma (Schermer et al., 1984). These collisions in the last 150 My may represent the beginnings of a new supercontinent (Condie, 1998). If they do, the breakup phase of Pangea and the growth phase of this new supercontinent significantly overlap in time (Fig. 9.1). During the last 500 My, Sr isotopes in marine carbonates have shown considerable variation (Fig. 9.3). Overall, they parallel the complex Gondwana–Pangea supercontinent history. The minima at 450 and 150 Ma may reflect the fragmentation of Laurasia (Laurentia–Baltica) and Pangea, respectively. The high isotopic ratios in the last 60 My probably reflect the collision of India with Asia and the uplift of the Himalayas (Harris, 1995). Other Sr isotopic peaks in the Paleozoic may reflect continental collisions in the assembly of Pangea, such as the Taconic orogeny in the Ordovician (400 Ma), the Acadian orogeny in the Devonian (360 Ma), the Hercynian orogeny (about 300 Ma), and the collision that produced the Ural Mountains (about 280 Ma). ... - tailieumienphi.vn