EARTH AS AN EVOLVING PLANETARY SYSTEM Part 2

Exhumation and Cratonization 33 Cratonization Although cratons have long been recognized as an important part of the continental crust, their origin and evolution is still not well understood. Most investigators agree that cratons are the end product of collisional orogenesis; thus, they are the building blocks of continents. Just how orogens evolve into cratons and how long it takes, however, is not well known. Although studies of collisional orogens show that most are characterized by clockwise P–T–t paths (Thompson and Ridley, 1987; Brown, 1993), the uplift–exhumation segments of the P–T paths are poorly constrained (Martignole, 1992). In terms of craton development, the less than 500° C portion of the P–T–t path is most important. Using a variety of radiogenic isotopic systems and estimated closure temperatures in various minerals, it is possible to track the cooling histories of crustal segments and, when coupled with thermobarometry, the uplift–exhumation histories. Results suggest wide variation in cooling and uplift rates; most orogens having cooling rates <2° C/My, whereas a few (such as southern Brittany) cool at rates <10° C/My (Fig. 2.11). In most cases, it would appear to take a minimum of 300 My to make a craton. Some terranes, such as Enderby Land in Antarctica, have had long, exceedingly complex cooling histo-ries lasting more than 2 Gy. Many orogens, such as the Grenville orogen in eastern Canada, have been exhumed as indicated by unconformably overlying sediments, reheated during subsequent burial, and then reexhumed (Heizler, 1993). In some instances, postorogenic thermal events such as plutonism and metamorphism have thermally overprinted earlier segments of an orogen’s cooling history such that only the very early high-temperature history (<500° C) and perhaps the latest exhumation record (<300° C) are preserved. Fission track ages suggest that final uplift and exhumation of some orogens, such as the 1.9-Ga Trans-Hudson orogen in central Canada, may be related to the early stages of supercontinent fragmentation. An important yet poorly understood aspect of cratonization is how terranes that amal-gamate during a continent–continent collision evolve into a craton. Does each terrane maintain its own identity and have its own cooling and uplift history? Or do terranes 900 800 Zircon/Garnet 700 600 Monazite Sphene Adirondack highlands 1050 Ma Southern brittany 400 Ma Pikwitonei 2640 Ma Taltson 2000 Ma S. India (Krishnagiri) 2500 Ma Figure 2.11 Cooling histories of several orogens. Ages of maximum temperatures are given in the explanation and equated to zero age on the cooling age axis. Blocking 500 Hornblende 400 300 Muscovite/Rutile Biotite temperature is the temperature at which the daughter isotope is trapped in a host mineral. Data from Harley and Black 200 Apatite 100 0 50 100 150 200 Cooling age (My) KFeldspar 250 300 350 (1987), Dallmeyer and Brown (1992), Mezger et. al. (1991) and Kontak and Reynolds (1994). 34 The Crust anneal to each other at an early stage so that the entire orogen cools and is elevated as a unit? What is the effect of widespread posttectonic plutonism? Does it overprint and erase important segments of the orogen cooling history? It is well known that crustal cooling curves are not always equivalent to exhumation curves (Thompson and Ridley, 1987). Some granulite-grade blocks appear to have undergone long periods of isobaric (constant depth) cooling before exhumation. Also, discrete thermal events can completely or partially reset thermochronometers without an obvious geologic rock record, and this can lead to erroneous conclusions regarding average cooling rates (Heizler, 1993). Posttectonic plutonism, which follows major deformation, or multiple deformation of an orogen can lead to a complex cooling history. Widespread posttectonic plutonism can perturb the cooling curve of a crustal segment, prolonging the cooling history (Fig. 2.12, left panel). In an even more complex scenario, a crustal domain can be exhumed, can be reburied as sediments accumulate in an overlying basin, can age for hundreds of millions of years around the same crustal level, and finally can be reexhumed (A in Fig. 2.12, right panel). In this example, all of the thermal history less than 400° C is lost by overprinting of the final thermal event. A second terrane, B, could be sutured to A during this event (S in Fig. 2.12, right panel), and both domains could be exhumed together. It is clear from these examples that much or all of a complex thermal history can be erased by the last thermal event, producing an apparent gap in the cratonization cooling curve. Processes in the Continental Crust Rheology The behavior of the continental crust under stress depends chiefly on the temperature and the duration of the stresses. The hotter the crust, the more it behaves like a ductile solid deforming by plastic flow. If it is cool, it behaves like an elastic solid deforming by brittle fracture and frictional gliding (Ranalli, 1991; Rutter and Brodie, 1992). The distribution of strength with depth in the crust varies with the tectonic setting, the strain rate, the thickness and composition of the crust, and the geotherm. The brittle–ductile transition corresponding to an average surface heat flow of 50 mW/m2 is around a 20-km depth, which corresponds to the depth limit of most shallow earthquakes. Even in the lower Figure 2.12 Two possible cooling scenarios in cratons. Left panel: Overprinting of a 800 Widespread plutonism 800 B posttectonic granite intrusion. Right panel: A 400 complex, multiple-event cooling history. S, suturing age of terranes A and B; Te, Ta,b final exhumation age. 0 Te Cooling age 400 A S 0 Ta,b Cooling age Process in the Continental Crust 35 crust, however, if stress is applied rapidly it may deform by fracture; likewise, if pore fluids are present in the upper crust—weakening it—and stresses are applied slowly, the upper crust may deform plastically. In regions of low heat flow, such as shields and plat-forms, brittle fracture may extend into the lower crust or even into the upper mantle because mafic and ultramafic rocks can be resistant to plastic failure at these depths; thus, brittle faulting is the only way they can deform. Lithologic changes at these depths, the most important of which is at the Moho, are also likely to be rheological discontinuities. Examples of two rheological profiles of the crust and subcontinental lithosphere are shown in Figure 2.13. The brittle–ductile transition occurs around a 20-km depth in the rift, whereas in the cooler and stronger Proterozoic shield, it occurs around 30 km. In both cases, the strength of the ductile lower crust decreases with increasing depth, reaching a minimum at the Moho. The rapid increase in strength beneath the Moho chiefly reflects the increase in olivine, which is stronger than pyroxenes and feldspars. The rheological base of the lithosphere, generally taken as a strength around 1 MPa, occurs 55 km beneath the rift and 120 km beneath the Proterozoic shield. In general, the brittle–ductile transition occurs at relatively shallow depths in warm and young crust (10–20 km), whereas in cool and old crust, it occurs at greater depths (20–30 km). Role of Fluids and Crustal Melts Fluid transport in the crust is an important process affecting both rheology and chemical evolution. Because crustal fluids are mostly inaccessible for direct observation, this process is poorly understood and difficult to study. Studies of fluid inclusions trapped in 0 20 Continental Rift 40 Figure 2.13 Rheological profile of the East African rift and the Proterozoic shield in East Africa. Strength expressed as the difference between maximum and minimum compressive stresses (s1 and s3, respectively). Diagram modified from Ranalli (1991). 60 80 Proterozoic Shield 100 −1.0 0 1.0 2.0 3.0 Strength log (s1-s3) (MPa) 36 The Crust metamorphic and igneous minerals indicate that shallow crustal fluids are chiefly water, whereas deep crustal fluids are mixtures of water and CO2, and both contain various dissolved species (Bohlen, 1991; Wickham, 1992). Fluids are reactive with silicate melts, and in the lower crust they can promote melting and can change the chemical and isotopic composition of rocks. In the lower crust, only small amounts of fluid can be generated by the breakdown of hydrous minerals such as biotite and hornblende. Hence, the only major source of fluids in the lower crust is the mantle. Studies of xenoliths suggest that the mantle lithosphere provides a potentially large source for CO2 in the lower crust, and the principal source for CO2 may be important in the production of deep crustal granulites. The formation of granitic melts in the lower crust and their transfer to shallower depths are fundamental processes leading the chemical differentiation of the continents. This is particularly important in arcs and collisional orogens. The melt-producing capacity of a source rock in the lower crust is determined chiefly by its chemical composition but also depends on temperature regime and fluid content (Brown et al., 1995). Orogens that include a large volume of juvenile volcanics and sediments are more fertile (high melt-producing capacity) than those that include chiefly older basement rocks from which fluids and melts have been extracted (Vielzeuf et al., 1990). A fertile lower crust can generate a range of granitic melt compositions and leave behind a residue of granulites. Segregation of melt from source rocks can occur by several processes, and just how much and how fast melt is segregated is not well known. These depend, however, on whether deformation occurs concurrently with melt segregation. Experiments indicate that melt segregation is enhanced by increased fluid pressures and fracturing of surrounding rocks. Modeling suggests that shear-induced compaction can drive melt into veins that transfer it rapidly to shallow crustal levels (Rutter and Neumann, 1995). Crustal Composition Approaches Several approaches have been used to estimate the chemical and mineralogical composi-tion of the crust. One of the earliest methods to estimate the composition of the upper continental crust was based on chemical analysis of glacial clays, which were assumed to be representative of the composition of large portions of the upper continental crust. Estimates of total continental composition were based on mixing average basalt and granite compositions in ratios generally ranging from 1:1 to 1:3 (Taylor and McLennan, 1985) or on weighting the compositions of various igneous, metamorphic, and sedimentary rocks according to their inferred abundances in the crust (Ronov and Yaroshevsky, 1969). Probably the most accurate estimates of the composition of the upper continental crust come from the extensive sampling of rocks exhumed from varying depths in Precambrian shields and from the composition of Phanerozoic shales (Taylor and McLennan, 1985; Condie, 1993). Because the lower continental crust is not accessible for sampling, indirect approaches must be used. These include (1) measuring seismic-wave velocities of crustal Crustal Composition 37 rocks in the laboratory at appropriate temperatures and pressures and comparing these with observed velocity distributions in the crust, (2) sampling and analyzing rocks from blocks of continental crust exhumed from middle to lower crustal depths, and (3) analyz-ing xenoliths of lower crustal rocks brought to the surface during volcanic eruptions. The composition of oceanic crust is estimated from the composition of rocks in ophiolites and from shallow cores into the sediment and basement layers of oceanic crust retrieved by the Ocean Drilling Project. Results are again constrained by seismic velocity distributions in the oceanic crust. Before describing the chemical composition of the crust, I will review the major sources of data. Seismic-Wave Velocities Because seismic-wave velocities are related to rock density and density is related to rock composition, the measurement of these velocities provides an important constraint on the composition of both the oceanic and the continental crust (Rudnick and Fountain, 1995). Poisson’s ratio, which is the ratio of P-wave to S-wave velocity, is more diagnostic of crustal composition than either P-wave or S-wave data alone (Zandt and Ammon, 1995) (Table 2.1). Figure 2.14 shows average compressional-wave velocities (at 600 MPa and 300° C) in a variety of crustal rocks. Velocities slower than 6.0 km/sec are limited to serpentinite, metagraywacke, andesite, quartzite and basalt. Many rocks of diverse origins have veloc-ities between 6.0 and 6.5 km/sec, including slates, granites, altered basalts, and felsic granulites. With the exception of marble and anorthosite, which are probably minor com-ponents in the crust based on exposed blocks of lower crust and xenoliths, most rocks with velocities from 6.5 to 7.0 km/sec are mafic in composition and include amphibolites and mafic granulites without garnet (Holbrook et al., 1992; Christensen and Mooney, 1995). Harzburgite Mafic garnet granulite Eclogite Anorthosite Amphibolite Greenstone basalt Mafic granulite Diorite Diabase Quartz mica schist Felsic granulite Tonalitic gneiss Granite/granodiorite Phyllite Basalt Granitic gneiss Quartzite Metagraywacke Serpentinite Dunite Garnet lherzolite Figure 2.14 Average compressional-wave velocities and standard deviations at 600 MPa (20-km depth equivalent) and 300° C (average heat flow) for major rock types. Data from Christensen and Mooney (1995). 5 5.5 6 6.5 7 7.5 8 8.5 9 Compressional wave velocity (km/sec) ... - tailieumienphi.vn