EARTH AS AN EVOLVING PLANETARY SYSTEM Part 2 potx

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 histories 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 Exhumation and Cratonization 33 100 200 300 400 500 600 700 800 900 1000 50 150 200 250 300 35 0 Blocking temperature (°C) Cooling age (My) Zircon/Garnet Hornblende Biotite Muscovite/Rutile KFeldspar Monazite Apatite 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 temperature is the temperature at which the daughter isotope is trapped in a host mineral. Data from Harley and Black (1987), Dallmeyer and Brown (1992), Mezger et. al. (1991) and Kontak and Reynolds (1994). 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/m 2 is around a 20-km depth, which corresponds to the depth limit of most shallow earthquakes. Even in the lower 34 The Crust Te Cooling age Blocking temperature °C 800 400 0 Ta,b Cooling age Blocking temperature °C 800 400 0 Widespread plutonism S B A Figure 2.12 Two possible cooling scenarios in cratons. Left panel: Overprinting of a posttectonic granite intrusion. Right panel: A complex, multiple-event cooling history. S, suturing age of terranes A and B; Te, Ta,b final exhumation age. 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 platforms, 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 Process in the Continental Crust 35 100 80 60 40 20 0 Depth (km) −1.0 0 1.0 2.0 3.0 Strength log (σ 1 -σ 3 ) (MPa) Proterozoic Shield Continental Rift LHERZOLITE GABBRO TONALITE 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 (σ 1 and σ 3 , respectively). Diagram modified from Ranalli (1991). metamorphic and igneous minerals indicate that shallow crustal fluids are chiefly water, whereas deep crustal fluids are mixtures of water and CO 2 , 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 CO 2 in the lower crust, and the principal source for CO 2 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 composition 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 36 The Crust 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) analyzing 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 velocities 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 components 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). Crustal Composition 37 5 5.5 6 6.5 7 7.5 8 8.5 9 Compressional wave velocity (km/sec) Dunite Eclogite Garnet lherzolite Mafic garnet granulite Gabbro Anorthosite Amphibolite Marble Mafic granulite Greenstone basalt Diabase Diorite Felsic granulite Quartz mica schist Granite/granodiorite Tonalitic gneiss Phyllite Slate Granitic gneiss Basalt Quartzite Andesite Metagraywacke Serpentinite Harzburgite 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). 38 The Crust Rocks with average velocities from 7.0 to 7.5 km/sec include gabbro and mafic garnet granulite, and velocities faster than 7.5 km/sec are limited to nonserpentinized ultramafic rocks and eclogite (a high-pressure mafic rock). It is important to note that the order of increasing velocity in Figure 2.14 is not a simple function of increasing metamorphic grade. For instance, low-, medium-, and high-grade metamorphic rocks all fall in the range from 6.0 to 7.5 km/sec. Although rock types in the upper continental crust are reasonably well known, the distribution of rock types in the lower crust remains uncertain. Platform lower crust, although it has relatively high S-wave velocities, shows similar Poisson’s ratios to collisional orogens (Fig. 2.15a; Tables 2.1 and 2.2). The lower crust of continental rifts, however, shows distinctly lower velocities, a feature that would appear to reflect hotter temperatures in the lower crust. Two observations are immediately apparent from the measured rock velocities summarized in Figures 2.14 and 2.15b: (1) the velocity distribution in the lower crust indicates compositional heterogeneity, and (2) metapelitic rocks overlap in velocity with mafic and felsic igneous and metamorphic rocks. It is also interesting that with the exception of rifts, mean lower crustal velocities are strikingly similar to mafic rock velocities. However, because of the overlap in velocities of rocks of different compositions and origins, it is not possible to assign unique rock compositions to the 5.0 (a) ( b ) 4.5 4.0 3.5 3.0 5.0 4.5 4.0 3.5 3.0 6.0 6.5 7.0 7.5 8.0 8.5 6.0 6.5 7.0 7.5 8.0 8.5 P wave velocity (km/sec) P wave velocit y( km/sec ) S wave velocity (km/sec) S wave velocity (km/sec) Rifts Metapelites Ultramafic rocks & eclogites 0.30 0.30 0.20 0.20 Shields & platforms Orogens Mafic rocks Felsic rocks Figure 2.15 Compressional-wave versus shear-wave velocity diagrams showing lower crust of various crustal types (a) and fields of various crustal and upper mantle rocks (b). Velocities are normalized to a 20-km depth and room temperature. Dashed lines are Poisson’s ratio (0.5{1–1/[(Vp/Vs) 2 –1]}). Modified from Rudnick and Fountain (1995). lower crust from seismic velocity data alone. When coupled with xenolith data, however, the seismic velocity distributions suggest that the lower continental crust is composed largely of mafic granulites, gabbros, and amphibolites (50–65%), with up to 10% metapelite, and that the remainder is intermediate to felsic granulite (Rudnick and Fountain, 1995). Based on seismic data, however, the lower crust in the Archean Kaapvaal craton in southern Africa appears to be felsic to intermediate in composition (James et al., 2003). In common rock types, Poisson’s ratio (σ) varies from about 0.20 to 0.35 and is particularly sensitive to composition. Increasing silica content lowers s, and increasing Fe and Mg increases it (Zandt and Ammon, 1995). The average value of s in the continental crust shows a good correlation with crustal type (Fig. 2.16; Table 2.1). Precambrian shield s values are consistently high, averaging 0.29, and platforms average about 0.27. The lower s in platforms and Paleozoic orogens appears to reflect the silica-rich sediments that add 4 to 5 km of crustal thickness to the average shield (Table 2.1). In Meso–Cenozoic orogens, however, s is even lower but more variable, reflecting some combination of lithologic and thermal differences in the young orogenic crust. The high ratios in continental- margin arcs may reflect the importance of mafic rocks in the root zones of these arcs, although again the variation in s is significant. The origin of the Moho continues to be a subject of widespread interest (Jarchow and Thompson, 1989). Because the oceanic Moho is exposed in many ophiolites, it is better known than the continental Moho. From seismic velocity distributions and from ophio- lite studies, the oceanic Moho is probably a complex transition zone from 0 to 3 km thick and between mixed mafic and ultramafic igneous cumulates in the crust and the harzbur- gites (orthopyroxene–olivine rocks) in the upper mantle. It would appear that large tectonic lenses of differing lithologies occur at the oceanic Moho and that these are the products of ductile deformation along the boundary. The continental Moho is considerably more complex and varies in nature with crustal type and age (Griffin and O’Reilly, 1987). Experimental, geophysical, and xenolith data, however, do not favor a gabbro–eclogite transition to explain the continental Moho. Also, the absence of a correlation between surface heat flow and crustal thickness does not favor a garnet granulite–eclogite phase change Crustal Composition 39 Mesozoic-cenozoic Collisional orogens Paleozoic collisional Orogens Platforms Shields Continental-margin arcs Poisson's ratio 0.2 0.24 0.28 0.32 0.36 Figure 2.16 Mean values and one standard deviation of Poisson’s ratio σ in various crustal types. Data from Zandt and Ammon (1995). at the Moho. Beneath platforms and shields, the Moho is only weakly (or not at all) reflec- tive, suggesting the existence of a relatively thick transition zone (<3 km) composed of mixed mafic granulites, eclogites, and lherzolites with no strong reflecting surfaces. Seismic Reflections in the Lower Continental Crust Many explanations have been suggested for the strong seismic reflectors found in much of the lower continental crust (Nelson, 1991; Mooney and Meissner, 1992) (Fig. 2.17). The most likely causes fall into one of three categories: 1. Layers in which fluids are concentrated 2. Strain banding developed from ductile deformation 3. Lithologic layering The jury is still out on the role of deep crustal fluids. Some investigators argue that phys- ical conditions in the lower crust allow up to 4% of saline pore waters and that the high electrical conductivity of the lower crust supports such a model. In this model, the seismic reflections are produced by layers with strong porosity contrasts. However, textural and mineralogical evidence from deep crustal rocks exposed at the surface and from xenoliths do not have high porosities, thus contradicting this idea. Deformation bands hold more promise, at least for some of the lower crustal reflections, in that shear zones exposed at the surface can be traced to known seismic reflectors at depth in shallow crust (Mooney and Meissner, 1992). Some lower crustal reflection patterns in Precambrian cratons preserve structures that date from ancient collisional events, such as in the Trans-Hudson orogen of Paleoproterozoic age in Canada. At least in these cases, it would appear that the reflections are caused by tectonic boundaries or by syntectonic igneous intrusions. In extended crust, such as that found in rifts, ductile shearing in the lower crust may enhance metamorphic or igneous layering. 40 The Crust Seismic section Line drawing Histogram 20 km 20 km 0 6 12 0 6 12 0 6 12 time/s time/s T o p o f r e f l e c t i v e l o w e r c r u s t Reflection moho Figure 2.17 Seismic reflection profile southwest of England. Also shown is a line drawing of the data and a histogram of reflection depths summed over 1-sec travel-time intervals. Modified from Klemperer (1987). The probable cause of many lower crustal reflections is lithologic layering, caused by mafic sills; compositional layering in mafic intrusions; or metamorphic fabrics. Supporting this conclusion are some shallow reflectors in the crust, which have been traced to the surface, that are caused by such layering (Percival et al., 1989). Furthermore, a bimodal distribution of acoustic impedance in the lower crust favors layered sequences of rocks, especially interlayered mafic and felsic units (Goff et al., 1994). Also, models of reflectivity in the Ivrea zone (a fragment of mafic lower crust faulted to the surface in the Alps) show that lower crustal reflections are expected when mafic rocks are interlayered with felsic rocks (Holliger et al., 1993). Seismic reflectivity in the lower crust is widespread and occurs in crustal types with differing heat-flow characteristics, favoring a single origin for most reflectors. From studies of exhumed lower crust and lower crustal xenoliths, it would seem that most lower crustal reflections are caused by mafic intrusions and, in some instances, that the reflections have been enhanced by later ductile deformation. Sampling of Precambrian Shields Widespread sampling of metamorphic terranes exposed in Precambrian shields and especially in the Canadian shield has provided an extensive sample base to estimate both the chemical and the lithologic composition of the upper part of the Precambrian continental crust (Shaw et al., 1986; Condie, 1993). Both individual and composite samples have been analyzed. Results indicate that although the upper crust is lithologically hetero- geneous, granitoids of granodiorite to tonalite composition dominate and the weighted average composition is that of granodiorite. Use of Fine-Grained Terrigenous Sediments Fine-grained terrigenous sediments may represent well-mixed samples of the upper continental crust and thus provide a means of estimating upper crustal composition (Taylor and McLennan, 1985). However, to use sediments to estimate crustal composition, it is necessary to evaluate losses and gains of elements during weathering, erosion, deposition, and diagenesis. Elements, such as rare earth elements (REE), Th, and Sc that are relatively insoluble in natural waters and have short residence times in seawater (<10 3 y) may be transferred almost totally into terrigenous clastic sediments. The remarkable uniformity of REE in pelites and loess compared with the great variability observed in igneous source rocks attests to the efficiency of mixing during erosion and deposition. Studies of REE and element ratios such as La/Sc, La/Yb, and Cr/Th indicate that they remain relatively unaffected by weathering and diagenesis. REE distributions are especially constant in shales and resemble REE distributions in weighted averages from Precambrian shields. With some notable exceptions (Condie, 1993), estimates of the average composition of the upper continental crust using the composition of shales are in remarkable agreement with the weighted chemical averages determined from rocks exposed in Precambrian shields. Crustal Composition 41 Exhumed Crustal Blocks Several blocks of middle to lower continental crust have been recognized in Precambrian shields or collisional orogens, the best known of which are the Kapuskasing uplift in southern Canada (Percival et al., 1992; Percival and West, 1994) and the Ivrea Complex in Italy (Sinigoi et al., 1994). Four mechanisms have been suggested to bring these deep crustal sections to the surface: (1) large thrust sheets formed during continent–continent collisions, (2) transpressional faulting, (3) broad tilting of a large segment of crust, and (4) asteroid impacts. However, tectonic settings at the times of formation of rocks within the uplifted blocks appear to be collisional orogens, island arcs, or continental rifts. Common to all studied sections are high-grade metamorphic rocks that formed at depths from 20 to 25 km with a few, such as the Kohistan arc in Pakistan, coming from depths as great as 40 to 50 km. Metamorphic temperatures recorded in the blocks are typically in the range from 700 to 850° C. All blocks consist chiefly of felsic components at shallow structural levels and mixed mafic, intermediate, and felsic components at deeper levels. Commonly, lithologic and metamorphic features in uplifted blocks are persistent over lateral distances greater than 1000 km, as evidenced by three deep crustal exposures in the Superior province in southern Canada (Percival et al., 1992). Examples of five sections of middle to lower continental crust are given in Figure 2.18. Each section is a schematic illustrating the relative abundances of major rock types, and the base of each section is a major thrust fault. The greatest depths exposed in each section are 25 to 35 km. Each column has a lower granulite zone, with mafic granulites dominating in three sections and felsic granulites in the other two. The sections show considerable 42 The Crust Volcanics and sediments Granite Felsic gneisses Amphibolite Anorthosite Mafic-ultramafic bodies Felsic granulites Mafic granulites Ivrea zone Fraser range Musgrave range Pikwitonei belt Kasila series Figure 2.18 Generalized cross-sections of continental crust based on exhumed sections of deep crustal rocks. Modified from Fountain and Salisbury (1981). [...]... Ni Middle 66.3 0.7 14.9 4.68 2. 46 0.07 3.55 3.43 2. 85 0. 12 87 26 9 626 9.1 18 2. 4 1 62 4.4 10.3 0. 82 25 29 59.4 4.83 1.05 2. 02 86 1 12 18 60 60.6 0.8 15.5 6.4 3.4 0.1 5.1 3 .2 2.0 0.1 62 281 4 02 6.1 15.3 1.6 125 4.0 8 0.6 22 17 45 4.4 1.5 2. 3 118 150 25 70 Lower 52. 3 0.54 16.6 8.4 7.1 0.1 9.4 2. 6 0.6 0.1 11 348 25 9 1 .2 4 .2 0 .2 68 1.9 5 0.6 16 8 20 2. 8 1.1 1.5 196 21 5 38 88 Oceanic Crust Total 59.7 0.68 15.7... 3.1 1.8 0.11 53 29 9 429 5.5 13 1.4 118 3.4 7.8 0.7 21 18 42 4.0 1 .2 1.9 133 159 27 73 50.5 1.6 15.3 10.4 7.6 0 .2 11.3 2. 7 0 .2 0 .2 1 90 7 0.1 0.3 0.05 74 2. 1 2. 3 0.13 28 2. 5 7.5 2. 6 1.0 3.1 27 5 25 0 47 150 Major elements in weight percentage of the oxide and trace elements in ppm (parts per million) Lower–middle crust from Rudnick and Fountain (1995), upper crust from Condie (1993), and oceanic crust (NMORB)... orogens composed of terranes (Karlstrom et al., 20 01) The Alps, Himalayas, and American Cordillera are Phanerozoic examples of orogens composed of terranes (Fig 2. 22) Most crustal provinces and orogens are composed of terranes, and in turn, cratons are composed of exhumed orogens You might consider terranes as the basic building blocks of continents and terrane collision as a major means by which continents... deformation and high-grade metamorphism in which new metamorphic zircons formed in the gneiss Grain 20 has a slightly 0.8 3600 3453±8 Trondhjemitic Gneiss AGC 185 3400 20 6Pb /23 8U 0.7 0.6 Zircons in banded gneiss Zircons in leucocratic vein 4 3166±4 3000 0.5 3505 24 320 0 2 1 25 00 3 0.4 20 00 2 4 3 1 5 0.3 1500 Grain 6 0 .2 Grain 20 6 0.1 0.1 mm 20 7Pb /23 5U Grain 4 0.0 0 5 10 15 20 25 30 35 Figure 2. 23 U-Pb... Eruption rates of flood basalts and oceanic-plateau basalts range from about 0.5 to less than 1 km3/year, considerably greater than rates typical of ocean ridges or volcanic islands such as Hawaii (with rates of 0. 02 0.05 km3/year) (Carlson, 1991; White and McKenzie, 1995) Continental flood basalts tend to be lower in Mg and Fe and other compatible elements than MORB or island basalts They typically show... ion probe analyses of zircons from a trondhjemitic gneiss in northeastern Swaziland, southern Africa Concordia is the bold solid line defined by concordant 20 6Pb /23 8U and 20 7Pb /23 5U ages From Kroner et al (1989) 51 52 The Crust discordant age of 3166 ± 4 Ma and comes from a later granitic vein that crosses the rock Other discordant data points in Figure 2. 23 cannot be fit to regression lines and reflect... volcanics related to mantle plumes (Condie, 20 01) About 10% of the ocean floors are covered by oceanic plateaus and aseismic ridges, and more than 100 are known (Fig 3.7), many of which are in the western Pacific (Coffin and Eldholm, 1994) These features rise thousands of meters above the seafloor, and some, such as the Seychelles Bank in the Indian Ocean (Fig 3.7), rise above sea level Some have granitic... pillow basalts and hyaloclastic breccias comprise most of the succession with only a thin capping of pelagic limestone, shale, and chert Some oceanic plateaus (like Kerguelen in the southern Indian Ocean) emerge above sea level and are capped with subaerial basalt flows and associated pyroclastic volcanics Oceanic-plateau basalts are largely tholeiites with only minor amounts of alkali basalt, and most... al., 20 03) Numerous potential terranes exist in the oceans today and are particularly abundant in the Pacific basin (Fig 2. 21) Continental crust may be fragmented and dispersed by rifting or strike-slip faulting In western North America, dispersion is occurring along transform faults such as the San Andreas and Fairweather, and in New Zealand movement along the Alpine transform fault is fragmenting the... supercontinent is Pangea, which formed between 450 and 320 Ma and includes most of the existing continents (Fig 2. 26) Pangea began to fragment about 160 Ma and is still dispersing today Gondwana is a Southern Hemisphere supercontinent composed principally of South America, Africa, Arabia, Madagascar, India, Antarctica, and Australia (Fig 2. 27) It formed in the latest Neoproterozoic and was largely completed . 3 .2 2.6 3.1 2. 7 K 2 O 2. 85 2. 0 0.6 1.8 0 .2 P 2 O 5 0. 12 0.1 0.1 0.11 0 .2 Rb 87 62 11 53 1 Sr 26 9 28 1 348 29 9 90 Ba 626 4 02 259 429 7 Th 9.1 6.1 1 .2 5.5 0.1 Pb 18 15.3 4 .2 13 0.3 U 2. 4 1.6 0 .2. 1 62 125 68 118 74 Hf 4.4 4.0 1.9 3.4 2. 1 Nb 10.3 8 5 7.8 2. 3 Ta 0. 82 0.6 0.6 0.7 0.13 Y25 22 1 621 28 La 29 17 8 18 2. 5 Ce 59.4 45 20 42 7.5 Sm 4.83 4.4 2. 8 4.0 2. 6 Eu 1.05 1.5 1.1 1 .2 1.0 Yb 2. 02. (1989). 0 5 10 15 20 25 30 35 0.8 0.7 0.6 0.5 0.4 0.3 0 .2 0.1 0.0 Grain 4 0.1 mm Grain 20 Grain 6 20 7 Pb/ 23 5 U 6 5 3 2 4 3 2 4 1 1 3505 24 3453±8 3600 3400 320 0 3000 25 00 20 00 1500 3166±4 Trondhjemitic

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