The generation and evolution of the continental crust

C.J. Hawkesworth; B. Dhuime; A.B. Pietranik; P.A. Cawood; A.I.S. Kemp; C.D. Storey

doi:10.1144/0016-76492009-072

Abstract:

The continental crust is the archive of the geological history of the Earth. Only 7% of the crust is older than 2.5 Ga, and yet significantly more crust was generated before 2.5 Ga than subsequently. Zircons offer robust records of the magmatic and crust-forming events preserved in the continental crust. They yield marked peaks of ages of crystallization and of crust formation. The latter might reflect periods of high rates of crust generation, and as such be due to magmatism associated with deep-seated mantle plumes. Alternatively the peaks are artefacts of preservation, they mark the times of supercontinent formation, and magmas generated in some tectonic settings may be preferentially preserved. There is increasing evidence that depletion of the upper mantle was in response to early planetary differentiation events. Arguments in favour of large volumes of continental crust before the end of the Archaean, and the thickness of felsic and mafic crust, therefore rely on thermal models for the progressively cooling Earth. They are consistent with recent estimates that the rates of crust generation and destruction along modern subduction zones are strikingly similar. The implication is that the present volume of continental crust was established 2–3 Ga ago.

The continental crust constitutes some 40% of the surface area of the Earth, and yet it constitutes almost 70% of the total volume of the Earth's crust. It is andesitic in composition, 25–70 km thick, and less dense than the thinner (<10 km) oceanic crust of largely mafic composition. The differentiation of the crust of the Earth into these contrasting chemical–mechanical components in part reflects the horizontal movement of the lithosphere (crust and upper mantle) through plate tectonics. The contrasting density structure of continental and oceanic crust results in a pronounced bimodal elevation, a buoyant continental crust and an oceanic crust that is, except for very young rocks, gravitationally unstable (e.g. Cloos 1993) and sinks back into the asthenospheric mantle at subduction zones, resulting in no oceanic crust being older than 200 Ma.

The continental crust is the geological archive of Earth history, it is different from analogues on nearby planets, and it influences global climate by being a sink for CO₂ (e.g. Garrels & Perry 1974; Zhang & Zindler 1993; Lowe & Tice 2004). It is andesitic in composition, and such magmas are not commonly in equilibrium with the upper mantle (e.g. Rudnick 1995; Walter 2003). Most models for the generation of new continental crust therefore involve the generation of basalt and subsequent differentiation, by fractional crystallization and/or remelting, to higher silica compositions (Kuno 1968; Ellam & Hawkesworth 1988; Arndt & Goldstein 1989; Kay & Kay 1991; Rudnick 1995; Arculus 1999; Kemp & Hawkesworth 2003; Zandt et al. 2004; Plank 2005; Hawkesworth & Kemp 2006a,b) and return of residue or cumulate to the mantle. Differentiation of the continental crust primarily involves igneous processes, and an idealized crustal section consists of a lower part dominated by residue or cumulate and/or new mafic crust and an upper part composed mainly of rocks of granitic to granodioritic composition. The residence times of elements in the upper crust appear to be much longer than those in the lower crust (Hawkesworth & Kemp 2006b).

The longstanding questions are when the continental crust was generated, and how the processes involved in the generation of the continental crust have changed with time. This review is written at a time of considerable conceptual upheaval when old approaches are being questioned and new analytical techniques have recently come into play. It has long been argued that the upper mantle was depleted by the extraction of the continental crust, and so the two reservoirs are complementary (Jacobsen & Wasserburg 1979; O'Nions et al. 1980; Allègre et al. 1983). The implication was that the record of continental crust generation could then be broadly investigated from the evolution of the depleted upper mantle. If that depleted mantle reservoir is well mixed, its radiogenic isotope ratios should offer a more robust indication of the volumes of crust extracted than the remaining vestiges of old continental crust. Yet recent evidence suggests that the upper mantle was initially depleted by processes much older than the preserved continental crust (Carlson & Boyet 2008; Tolstikhin & Kramers 2008). Similarly, it has been widely assumed that the geological record provides a representative record of the evolution of the continental crust, and yet now there is increasing evidence that it may be biased by the tectonic settings in which the rocks were generated (Hawkesworth et al. 2009).

In this review we explore current thinking on when and how new continental crust was generated, and the extent to which there were peaks of crust generation and variations in the rates at which new crust was generated. We consider the constraints on the composition of early continental crust in the Hadean (older than 4.0 Ga), and links between granite magmatism, crustal growth and high-grade metamorphic events. It is timely also to review the approaches to these longstanding issues.

Old problems and new ways forward

The oldest known rocks on Earth are the 4 Ga Acasta Gneisses from the Slave Province in NW Canada (Bowring & Williams 1999), and the oldest minerals are detrital zircons that are up to 4.4 Ga old from the Jack Hills area in Western Australia (Wilde et al. 2001). The Jack Hills zircons, and their diverse inclusions of minerals such as quartz, feldspar, muscovite and monazite, suggest that geochemically evolved rocks, and hence by implication continental crust, were present from at least 4.4 Ga. However, the volume of continental crust cannot be assessed from the small number of isolated records that are preserved.

Figure 1 summarizes some of the data and interpretations that have been a focus for many discussions of the generation and evolution of the continental crust in the last 20 years. A number of models sought to describe the increase in the volume of continental crust through time. Many of the models are based on radiogenic isotopes; for example, the Nd isotope ratios of shales (McCulloch & Wasserburg 1978; O'Nions et al. 1983; Allègre & Rousseau 1984). Strictly speaking, they therefore describe the increase in the volume of continental crust that survived for long enough for the isotope ratios to change in response to radioactive decay, which is of the order of several hundred million years. As pointed out by Armstrong (1981), crust that is generated one day and destroyed soon thereafter will not be visible to records based on the radiogenic isotope ratios of crustal rocks. Formally, therefore, the curves, which are often described as showing increases in the volume of continental crust through time, in practice describe the volume of (relatively) stable continental crust. In general these curves are smooth (Fig. 1), and they contrast with the spiky distribution of the ages of igneous rocks derived from the mantle that therefore represent the generation of new continental crust (McCulloch & Bennett 1994; Condie 1998, 2005).

Fig. 1.

A histogram of the volume distribution of juvenile continental crust based on a compilation of U–Pb zircon ages integrated with Nd isotope ratios and lithological associations. This is compared with models of continental growth with 100% representing the present-day cumulative volume of crust (adapted from Condie 2005; after Cawood et al. 2009). Early models, such as that by Hurley & Rand (1969), were based on the geographical distribution of Rb–Sr and K–Ar isotope ages, and these are likely to have been reset by younger orogenic events. Some models suggested rapid early growth of continental crust, and a slower rate of growth or even a decrease in continental volume through time (Fyfe 1978; Reymer & Schubert 1984; Armstrong 1991). Other models require a more linear growth (unlabelled) or rapid growth during the late Archaean followed by steady-state growth driven by island arc magmatism (Taylor & McLennan 1985).

The details of the igneous record of new crustal growth remain difficult to establish, and this plot is included because, albeit in slightly different versions, it has been a focus of discussion for the last 20 years. The distinctive feature of the igneous record is that there are two pronounced age peaks at 2.7 and 1.9 Ga, with some compilations suggesting a third at 1.2 Ga, and then none in the last 1 Ga. One interpretation is that the age peaks reflect periods of exceptional rates of crustal growth, that such peaks are difficult to ascribe to processes linked to plate tectonics, and they may therefore be evidence for periods of crustal growth linked to the emplacement of superplumes (e.g. Condie 2004). The lack of conspicuous age peaks in the last billion years might in turn reflect the period in which the generation of new continental crust was largely dominated by plate-tectonic processes. The underlying question is the extent to which these age distributions are a true record of the generation and evolution of the continental crust, or whether they are artefacts of preservation in the geological record (e.g. Gurnis & Davies 1985, 1986; Hawkesworth et al. 2009).

One of the marked developments in recent years is our ability to measure isotope and trace elements in situ by ion microprobe and laser ablation inductively coupled plasma mass spectromery (LA-ICP-MS). It is now much easier to characterize what is being analysed, and there have been exciting new developments in topics as diverse as the generation of granites, the records preserved in detrital and inherited zircons, and the dating of high-grade metamorphic processes. Recent studies of crustal evolution have relied on zircons, which typically yield high-precision U–Pb ages and so have been the cornerstone of most attempts to develop an accurate geological time scale. Zircons can be analysed in situ for U–Pb, Hf and O isotopes and trace elements. They are resilient under most crustal conditions and they survive metamorphism and prolonged sedimentary recycling. Thus they arguably provide more representative records, particularly from older geological terrains, than the rock record. However, most zircons crystallize from medium- to high-silica magmas and so they more readily offer insights into the generation and evolution of more evolved compositions in the crust than into the processes of crust generation and the evolution of the depleted upper mantle.

Terms

The literature on the generation and evolution of the continental crust and the mantle contains a number of terms that are widely, but perhaps not consistently used. It is therefore helpful to start with definitions of some of these terms as they can mean different things to different people, which may in turn lead to fundamentally different implications. Following Rudnick & Gao (2003), the continental crust is taken to extend vertically from the Earth's surface to the Mohorovičić discontinuity and laterally to the break in slope in the continental shelf. It is the layer of granitic, sedimentary and metamorphic rocks that form the continents and the areas of shallow seabed close to their shores, the continental shelves. It is linked to the upper mantle and together these form the lithosphere, the rigid, highly viscous lid of the Earth that is divisible into a series of plates. The lithospheric plates range in thickness from less than 100 km for oceanic lithosphere, and probably as low as 10 km close to mid-ocean ridges, to up to 200–300 km under old Archaean nuclei. The plates move both with respect to each other along plate boundaries, and with respect to the underlying mantle asthenosphere from which they are separated by the low-velocity zone.

We are here concerned with the continental crust as a reservoir of rocks of relatively differentiated compositions at the surface of the Earth, and the volume that may be present at different times. We are interested in generic models for how the average ‘andesitic' composition of the continental crust may have been generated, and less with the details of how it may have been generated in different ways at different localities. New continental crust is generated from the mantle, but its average composition would not be in equilibrium with typical mantle compositions. Thus it is likely to have been generated in more than one stage as, for example, in the generation, crystallization and remelting of basalt.

The term crust generation is therefore used to denote the formation of new continental crust in a generic sense; that is, the emplacement of new magma directly from the mantle. The growth of continental crust is then formally the increase in its volume through time. This necessarily takes account both of the volumes of new crust generated and the amounts destroyed by erosion and returned to the mantle. In practice, growth of continental crust is difficult to tie down, because radiogenic isotopes constrain only the volume of crust that has been stable for long enough for significant differences in isotope ratios to be developed from radioactive decay. However, even short-lived crust may leave a legacy in the complementary depletion of the upper mantle. The assembly of continental crust from different segments that were generated elsewhere and juxtaposed tectonically increases the volume of continental crust in the region being considered, but not the volume of continental crust overall, in the sense that the assembled fragments were already present elsewhere. The proportion of continental crust is inversely proportional to oceanic and transitional crust on a constant radius Earth. Thus as the volume of continental crust has grown (assuming it has) then the proportion of oceanic crust has decreased.

In this review crustal recycling is taken to mean the recycling of continental crust back through the mantle and into the crust again. Crustal reworking is the remobilization of pre-existing crust by partial melting and/or erosion and sedimentation, but all at sites within the continental crust.

The timing of events: crystallization and crust formation ages

Accurate and precise ages underpin our understanding of the generation and the evolution of the continental crust. In practice, two types of ages are involved. The first is the geological age of the material being analysed, which can be the age of deposition of a sediment or the crystallization of a single mineral or an igneous rock. Zircons are widely used because they yield high-precision U–Pb ages, and these ages may be determined by dissolving the whole crystal, and by in situ isotope measurements using secondary ionization mass spectrometry (SIMS) or laser ablation ICP-MS. Zircons tend to yield relatively reliable isotope results, and they are arguably easier to interpret than whole-rock data, which may be more readily disturbed by metamorphic and low-temperature alteration events. The advantage of in situ measurements is that the material being analysed can be well characterized by microbeam imaging, in contrast to whole-rock analyses (Fig. 2). Moreover, small portions, 10–50 μm across, can be dated and so the history of more complicated grains can be unravelled.

Fig. 2.

(a) Cenozoic granulite-facies gneiss from the Hidaka Metamorphic belt, northern Japan. The coin is about 2 cm in diamter. (b) A cathodoluminesence image from a zircon recording crystallization events at 54 and 19 Ma (Kemp et al. 2007b). Leucosomes have developed with a consistent orientation, typically containing garnet. Granulite generation involved the reworking of older protoliths through enhanced heat flux associated with subduction zone retreat.

The second types of ages are those that constrain when the crustal source of igneous or sedimentary rock was itself derived from the mantle. These are critical to this discussion because they seek to constrain when new crust was generated, they are termed model ages, and they are widely used for Hf and Nd isotopes. The key assumption of these model ages is that the ratio of the parent to the daughter isotope (Lu/Hf and Sm/Nd) changes in the processes involved in the generation of new crust, and then it is not fractionated further by processes of remelting, erosion and sedimentation within the continental crust. Thus the parent/daughter ratio of the crustal source material can be estimated, and used to calculate the time when that portion of new crust was initially extracted from the mantle. These ages are known as model Nd and Hf ages, and they are best illustrated on isotope evolution diagrams of isotope ratios against time (Fig. 3). In many cases the Nd and Hf isotope ratios of sediments will be hybrid values averaging the contributions of material from different source rocks. They are less likely to reflect discrete crust-forming events. Thus it is helpful to distinguish model age that are derived from sedimentary sources from those derived from igneous sources, and this can be done using O isotopes.

Fig. 3.

Model Hf ages. A plot illustrating the changes in Hf isotope ratios with time in the bulk Earth, the depleted mantle and in average continental crust. The present-day Hf isotope ratio of a zircon is similar to that at the time it crystallized, because it has a very low Lu/Hf ratio. Its model Hf age is calculated as the age at which the evolution of the source of the magma from which the zircon crystallized intersects the curve for the depleted mantle. The (crustal) source of the magma is traditionally assumed to have the ¹⁷⁶Lu/¹⁷⁷Hf ratio of average continental crust.

The ¹⁸O/¹⁶O ratio, expressed as δ¹⁸O, is readily changed only by low-temperature and surficial processes, and so the δ¹⁸O of mantle-derived magmas (5.3 ± 0.3‰; Valley 2003) are different from those of rocks that have experienced a sedimentary cycle or low-temperature hydrothermal alteration on the sea floor, which have elevated δ¹⁸O (c. 7–25‰; Eiler 2001, and references therein). Magmas that contain a contribution from sedimentary rocks should therefore have elevated δ¹⁸O values, and these in turn will be preserved in zircons that crystallized from such magmas. Thus, high δ¹⁸O values in zircon are a ‘fingerprint' for a supracrustal component in the generation of felsic igneous rocks, with the implication that such zircons are likely to yield hybrid model ages. Empirical studies have established that oxygen diffusion in zircon is sufficiently sluggish that the original igneous δ¹⁸O value remains intact, even through protracted metamorphism and crustal fusion (King et al. 1998; Peck et al. 2003). O isotopes can be measured in situ in zircon with excellent precision (<0.5‰) by large radius ion microprobes that have multi-collector capability (Valley 2003), and so it is possible to determine U–Pb, Hf and O isotope and trace element data on the same grain of zircon (or part there of), all to high precision, and using in situ techniques.

Variations in Sm/Nd and Lu/Hf ratios and the composition of initial continental crust

Key points of discussion are the ages of crust generation, as inferred from model ages, and the composition of that initial crust. The calculation of model ages presumes that the parent/daughter ratios are constant in the dominant crustal lithologies. For Nd isotopes, it is generally accepted that the Sm/Nd ratios of different crustal rocks are similar (but see below), and so model Nd ages are routinely calculated using the measured Sm/Nd ratio of the sample analysed. In contrast, zircons are widely analysed for Hf isotopes because they have extremely low Lu/Hf ratios, and so their present-day Hf isotope ratios are similar to those when they crystallized. It follows that the Lu/Hf ratio of zircon is very different from that of the likely source rocks of their host magma in the continental crust, and so Hf isotopes in zircon model ages are typically calculated using the average Lu/Hf ratio of the continental crust (e.g. Griffin et al. 2002). A characteristic of the Hf isotope–time plot is that the slopes of straight lines depend on their parent–daughter Lu/Hf ratios. It follows that straight isotope evolution lines defined by analyses of zircons of different ages, but from magmas from similar source rocks, have slopes that reflect the Lu/Hf ratios of the crustal source rocks. In detail the Lu/Hf ratios of igneous rocks decrease with increasing silica and trace element indices of differentiation such as Rb/Sr (see Fig. 4, and the associated references); so if the slope of the Hf isotope evolution line for the crustal source rocks can be determined it should be possible to evaluate whether those crustal source rocks are mafic or granitic in composition. It is helpful, therefore, to evaluate the range of Lu/Hf and Sm/Nd ratios in common rock types.

Fig. 4.

Median Lu/Hf–SiO₂, Lu/Hf–Sm/Nd, Rb/Sr–SiO₂, Rb/Sr–Sm/Nd, and Rb/Sr–Lu/Hf in basalts and basaltic andesites [SiO₂ < 55%, mid-ocean ridge basalts (MORB): n = 331; oceanic island basalts (OIB): n = 1520; island arcs: n = 644; continental flood basalts (CFB): n = 2514], granitoids (SiO₂ ≥ 55%; n = 7371) and continental sediments (n = 63) (GEOROC compilation, http://georoc.mpch-mainz.gwdg.de/georoc/). Rb/Sr is used as an index of differentiation. Primitive mantle (PUM) and CI Chondrites after Sun & McDonough (1989), continental crust from Rudnick & Gao (2003), cratonic shales from Condie (1993) and GLOSS (global subducting sediment) from Plank & Langmuir (1998). The horizontal and vertical bars represent the deviation of the median from the mean and mode values reported in Table 1.

Figure 4 and Table 1 summarize the median values of the Sm/Nd and Lu/Hf ratios in basalts, granites and sedimentary rocks compiled from a database of c. 12 000 analyses from GEOROC (http://georoc.mpch-mainz.gwdg.de/georoc/). Overall there is a broad correlation between the two ratios but the range in the median values for Lu/Hf is almost twice that for Sm/Nd. Because the half-life of ¹⁷⁶Lu is a third that of ¹⁴⁷Sm, the parent/daughter ratio, and hence the bulk composition of early formed crust, is better approached using Hf rather than Nd isotopes. It follows that there is a greater range for Lu/Hf than for Sm/Nd for each rock type (Table 1). Island arc basalts and basaltic andesites, and mid-ocean ridge basalt (MORB), have higher Sm/Nd and Lu/Hf ratios than the rocks of the continental crust, and both ratios decrease from estimates of the average lower continental crust, through the values for the bulk crust to the upper crust. The abundances of insoluble elements in the upper crust are estimated from those in continental sediments, and so they both plot together in Figure 4. Granitic rocks and the bulk crust have similar Sm/Nd and Lu/Hf ratios, as do flood basalts and the lower crust. The average lower crustal composition is similar to estimates of model new continental crust; that is, the mantle-derived magma from which the more evolved compositions of the crust then differentiated (Hawkesworth & Kemp 2006a). However, the similarity in Sm/Nd and Lu/Hf ratios between flood basalts and the lower crust may be simply coincidental, and it should not be taken to imply that both were necessarily generated in similar settings. Current models attribute the trace element ratios of the continental crust to the combination of processes involved in the generation of new crust from the mantle, and to the subsequent foundering of lower crustal material (Ellam & Hawkesworth 1988; Arndt & Goldstein 1989; Kay & Kay 1991; Rudnick 1995; Arculus 1999; Kemp & Hawkesworth 2003; Zandt et al. 2004; Plank 2005; Hawkesworth & Kemp 2006a,b).

View this table:

Table 1

A summary of the SiO₂ contents and selected parent–daughter trace element ratios in selected geological units and rock types

In seeking to interpret the variations in Lu/Hf and Sm/Nd ratios it is helpful to evaluate how they vary with indices of magma differentiation. For crustal systems, Rb/Sr ratios are more sensitive indices of differentiation than bulk SiO₂ contents, and so Figure 4 summarizes Rb/Sr−SiO₂, Rb/Sr–Sm/Nd, and Rb/Sr–Lu/Hf for various rock units. A positive correlation is observed in the Rb/Sr–SiO₂ diagram and the Sm/Nd and Lu/Hf ratios show a negative correlation with Rb/Sr. What is striking is the shift in Lu/Hf and Sm/Nd from the arc basalts or basaltic andesites into the granitic rocks and sediments. In detail, therefore, Sm/Nd and Lu/Hf both decrease with increasing differentiation in the continental crust, and this may introduce additional uncertainties into the calculations of model ages, especially in strongly differentiated samples.

Zircons as archives of the continental crust

Zircons typically crystallize from granitic magmas with >60% SiO₂, although they are found in lower silica magmas in lesser abundance. They are therefore a feature of magmatic rocks in the upper continental crust, they yield high-precision (U–Th)–Pb crystallization ages, and they survive erosion and sedimentation to be preserved in continent-derived sedimentary rocks. Zircon has the unique combination of physiochemical resilience and high concentrations of important trace elements that include two radiogenic isotope systems of geochronological importance (namely, U–Pb, Th–Pb) and another (Lu–Hf) that is widely used as a tracer of crustal evolution. It follows that zircons offer an excellent record of the evolution of differentiated compositions in the continental crust, and it may be more difficult to use zircons to chart the generation of new crust and hence the evolution of the depleted mantle, for example. Continent-derived sediments, and in particular shales, have been widely used to obtain representative average compositions of the upper crust (e.g. Taylor & McLennan 1991; Condie 1993). This works well for insoluble elements, the REE, Y, Sc, Th and Nb, and also for Ti, Zr and Hf, but the abundances of the second group are controlled by the distribution of heavy minerals. Zircon is one such mineral, and so the question becomes how representative of their source terrain, and hence of the upper crust in a particular area, are zircons preserved in continental sediments. Zircons, for example, are concentrated in more mature sands, and yet average compositions, and Nd model ages, are estimated from shales.

Figure 5 summarizes U–Pb crystallization and Hf model ages on zircons from Australia in both magmatic rocks, and as detrital grains in sediments. Just over 17 000 zircons have been dated, and there are Hf model ages on over 2000 of them, and so they provide a good basis for exploring the information that is available from the study of zircons. The striking observation is that there are marked peaks in the ages of crystallization, although the peak at 2.7 Ga partly reflects the intensive study of economically important rocks of this age in Western Australia, and peaks of similar ages are present in both the magmatic and detrital zircons. It appears therefore that the detrital zircons offer representative records of the magmatic history seen in the exposed igneous provenance.

Fig. 5.

A summary of the U–Pb crystallization and Hf model ages on zircons from Australia both in magmatic rocks and as detrital minerals. This contrasts the peaks of crystallization ages for the zircons from magmatic rocks and as detrital grains in sediments with the broader distribution of the Hf model ages. The data sources are available upon request.

The peaks of crystallization ages primarily represent peaks of magmatic activity, although we shall return to the issue of whether the record may in some way be biased by the nature of the geological record. It is not possible to say from the crystallization ages whether these magmatic episodes involved the generation of significant volumes of new crust or not, but in principle this may be investigated by the determination of Hf isotope ratios, and hence of Hf model ages. In practice, two approaches have been used. First, the analysis of Hf and O isotopes in magmatic zircons (i.e. those that crystallized at the same time as their host rock) is used to unpick the relative contributions of mantle and crustal sources in their petrogenesis (Kemp et al. 2007a). Such detailed studies constrain the amounts of new crust generated in the different peaks of magmatic activity identified from the crystallization ages of the zircons.

Hf model ages are also measured in detrital and inherited zircons. However, at present there is little sense of the kinds of magma from which the zircons crystallized, apart from the observation that they tend to have high silica contents. The difficulty is that the magmas may contain mixed contributions from the mantle and the crust at the time they crystallized, and some magmas also contain contributions from sediments, which are in themselves typically derived from mixed sources. Thus, the Hf isotope record in the detrital and inherited zircons is likely to be dominated by processes of mixing within the crust, and so it is extremely difficult to establish the ages of the initial crust-forming events, except in cases where O isotopes are also available (Kemp et al. 2006; Pietranik et al. 2008). One consequence is that, taken together, the Hf model ages of the zircons from Australia show much less well-defined peaks than do the crystallization ages (Fig. 5). There are peaks of Hf model ages at c. 4.4 Ga, from the Jack Hills locality in Western Australia, and at c. 3.2 Ga, but they are much less marked subsequently. The implication is that these zircons provide a record that is dominated by crustal reworking, as is also seen for Nd isotopes in shales (Allègre & Rousseau 1984).

In practice, sedimentary and igneous rocks offer different records and perspectives on the evolution of the continental crust (see also Fig. 1). Igneous rocks are for the most part generated in magmatic provinces that are restricted in space and time, and the generation of new crust clearly involves magmatic processes. Sediments, in contrast, are mixtures of the material that was present in potential source regions at the time of deposition. Thus their Nd and Hf model ages are necessarily hybrid ages that in most cases do not reflect discrete crust-forming events; they are difficult to interpret because the ages, and the relative contributions of different source terrains, cannot be independently determined from whole-rock samples (e.g. Allègre & Rousseau 1984). Recently, two ways forward have been developed using zircons: (1) to use O isotopes to screen out zircons that crystallized from magmas that may contain a contribution from sedimentary source regions, and may therefore yield hybrid model ages; (2) to use zircons to constrain the source regions, and their relative contributions, represented in whole-rock sediments.

In subsequent sections we consider a series of case studies showing new advances and/or perspectives in a number of key areas pertaining to the continental growth and crust–mantle differentiation. These include the differentiation of the infant silicate Earth, the composition of nascent continents, rates of crust generation and preservation in the geological record, and the information available from the igneous and sedimentary records.

Differentiation of the infant silicate Earth: formation of early depleted and enriched reservoirs

The depleted mantle is traditionally regarded as the complementary reservoir to the continental crust (Jacobsen & Wasserburg 1979; O'Nions et al. 1980; Allègre et al. 1983). Isotopically it is depleted relative to the bulk Earth, it is sampled by MORB at the present day, and that depletion has typically been attributed to the extraction of the continental crust. In terms of radiogenic isotope ratios such arguments work well; the extraction of high Rb/Sr crust, for example, leaves behind a low Rb/Sr residual depleted mantle, and with time those reservoirs develop high and low ⁸⁷Sr/⁸⁶Sr ratios respectively. In detail, the processes involved are much debated. If the new material extracted from the mantle as a precursor to the continental crust is basaltic, the residual depleted mantle may be sufficiently infertile to generate more basalt in subsequent melting events, unless some process of refertilization is invoked. None the less, there has been considerable interest in determining the isotope evolution of the depleted mantle, both as a reference frame for the calculation of model ages (e.g. Fig. 3), and because it might constrain when the continental crust was generated.

The initial isotope ratios from whole-rock analyses of rocks derived from the mantle should in principle offer robust estimates of isotope ratios of the mantle at those times. Figure 6 compares models for the evolution of the depleted mantle based on initial whole-rock Hf isotope ratios with high ε_Hf values. The continuous line was calculated using the ¹⁷⁶Lu/¹⁷⁷Hf and ¹⁷⁶Hf/¹⁷⁷Hf ratios of the depleted mantle at the present day (Griffin et al. 2002) and this is the curve that is most widely used to calculate the model ages of crustal rocks. In practice, the observed maximum ε_Hf values do not simply increase with decreasing age, as they should in mantle with a constant Lu/Hf, and this has been attributed to enhanced crustal recycling in the Archaean (Bennett 2003). The implication is that the depleted mantle evolution would then be characterized by changes in its Lu/Hf through time (short-dashed line, Fig. 6). Critically, however, both these interpretations based on whole-rock compositions are different from those predicted from simple models in which the depleted mantle and the continental crust are treated as complementary reservoirs, and the crustal volume grows through time (the long-dashed line in Fig. 6 is the estimated ε_Hf evolution of the mantle adapted from ε_Nd in the mantle from Nägler & Kramers (1998)). Particularly for Archaean samples, it remains difficult to establish the extent to which the whole rocks have been disturbed isotopically, and some have higher initial Nd and Hf isotopes ratios than predicted by most models of early crust extraction. Thus there has been increasing interest in mineral archives that yield precise, typically U–Pb ages of crystallization, and preserve robust records of other isotope systems. Zircons have considerable potential, although because they tend to crystallize from relatively evolved magmas, they may offer greater insights into the evolution of the continental crust than the evolution of the depleted mantle.

Fig. 6.

Evolution of ε_Hf through time in the depleted mantle. The basalts, komatiites and Archaean TTG whole-rock compositions plotted are those with high ε_Hf and a range of ages. Continuous black line: depleted mantle evolution based on present-day values of ¹⁷⁶Lu/¹⁷⁷Hf = 0.0384 and ¹⁷⁶Hf/¹⁷⁷Hf = 0.28325 (Griffin et al. 2002). Short-dashed line: evolution with changing Lu/Hf through time based on maximum ε_Hf in whole rocks (Bennett 2003). Long-dashed line: evolution adapted from the model for the depleted mantle based on ε_Nd from Nägler & Kramers (1998), and the correlation between Nd and Hf isotopes after Vervoort et al. (1999). ε_Hf was calculated using present-day CHUR values of ¹⁷⁶Lu/¹⁷⁷Hf = 0.0336 and ¹⁷⁶Hf/¹⁷⁷Hf = 0.282785 (Bouvier et al. 2008).

Figure 7 contrasts laser ablation zircon Hf isotope data and solution Hf isotope data from digestion of zircon fractions from the Amîtsoq gneisses of southern West Greenland (Vervoort et al. 1996; Vervoort & Blichert-Toft 1999; Kemp et al. 2009a). It is striking that the zircons analysed from solution tend to have higher initial Hf isotope ratios, and therefore apparently offer more insight into the Hf isotope ratios of the depleted mantle. In contrast, positive ε_Hf values are not evident within the laser ablation zircon data, apart from a single analysis of a metamorphic overgrowth. The laser ablation Hf isotope data define a simple array that overlaps the less radiogenic end of the solution Hf isotope field. The slope of the array is consistent with the evolution of a crustal reservoir with ¹⁷⁶Lu/¹⁷⁷Hf c. 0.019, which was remelted at c. 3.81, 3.71 and 3.65 Ga to yield zircons of these ages, and such a ¹⁷⁶Lu/¹⁷⁷Hf ratio is consistent with mafic to intermediate crust (see Fig. 4). The array provides permissive evidence for extraction of the crustal source material from either a chondritic reservoir at c. 3.85 Ga or model depleted mantle at 4.0 Ga. There is no evidence from these laser ablation data for the addition of juvenile material between 3.81 and 3.65 Ga, which would generate vertical mixing arrays extending to higher ε_Hf.

Fig. 7.

A plot of ε_Hf v. inferred crystallization age for zircons from Greenland gneisses, comparing data obtained by solution analysis (Vervoort et al. 1996; Vervoort & Blichert-Toft 1999) with those measured by laser ablation using both standard Hf isotope and concurrent Pb–Hf isotope routines (ε_Hf values of the latter are calculated using the laser ablation ²⁰⁷Pb/²⁰⁶Pb age; error bars are indicated at 2 SE) (Kemp et al. 2009a). The short-dashed line depicts the evolution of a putative crustal reservoir (¹⁷⁶Hf/¹⁷⁷Hf = 0.019) derived from depleted mantle at 4.0 Ga (the depleted mantle evolution curve was calculated using mean present-day values of ¹⁷⁶Lu/¹⁷⁷Hf = 0.0384 and ¹⁷⁶Hf/¹⁷⁷Hf = 0.28325, after Griffin et al. 2002). The solution ε_Hf data from zircons of Amîtsoq gneiss GGU-110999 recalculate to −1.4 and −1.7 at the inferred magmatic age of 3.65 Ga, as shown by the fine arrowed line. The second arrowed line connects the ε_Hf values of the metamorphic zircon rim in GGU 125540 calculated at 3.65 Ga and 3.5 Ga. These highlight the sensitivity of ε_Hf values calculated for ancient zircons to crystallization age.

Of concern is how to reconcile the laser ablation data with the much larger solution Hf isotope datasets from the Amîtsoq Gneisses, in which positive ε_Hf values are prevalent (Fig. 7). Assuming that both datasets are representative, these data highlight the difficulty of interpretation of solution ε_Hf values from the Amîtsoq gneiss, given the complex zircon microstructures and U–Pb age dispersions (Whitehouse et al. 1999). The most appropriate age at which to calculate the ε_Hf values for the Amîtsoq Gneiss zircons requires knowledge of the proportion of Hf contributed from the different growth zones. This is not easy to determine for solution analysis, as bulk zircon digestion homogenizes Hf derived from domains of different age, and potentially, of different ¹⁷⁶Hf/¹⁷⁷Hf. However, they can be resolved by in situ U–Pb and Lu–Hf analysis of zircons from the same samples from which the solution data were acquired.

The issues raised in the discussion of Figure 7 are important in the study of the older, more complex Archaean terrains, and in the analysis of complex detrital zircons, but they are less significant for the analysis of simple magmatic zircons. Pietranik et al. (2009) explored how the large volumes of Hf isotope data on concordant zircons might now be used to constrain the evolution of the depleted mantle. Figure 8 illustrates that only a small number of zircons plot on the depleted mantle evolution line of Griffin et al. (2002) and these are all <2.7 Ga. Many zircons also plot above the curve that reflects the sigmoidal increase in the volume of continental crust through time (the curve modified after Nägler & Kramers (1998) in Fig. 6). The implications are as follows.

Fig. 8.

A zircon-based model for the evolution of ε_Hf in the depleted mantle. The bold continuous line is a regression through open circles representing depleted mantle compositions based on the approach of Pietranik et al. (2009). The short-dashed line is the two-stage evolution of the depleted mantle based on Tolstikhin et al. (2006); the mantle was depleted immediately after accretion (c. 4.6 Ga) by extraction of an early enriched reservoir, and the crust was extracted later following the crustal growth rates of Taylor & McLennan (1995); 90% of the mantle was depleted. The fine grey line shows the depleted mantle evolution based on the present-day values for the depleted mantle of ¹⁷⁶Hf/¹⁷⁷Hf = 0.28325 and ¹⁷⁶Lu/¹⁷⁷Hf = 0.0384 (Griffin et al. 2002). The long-dashed curve and the CHUR parameters are as in Figure 6.

(1) The depleted mantle evolution line of Griffin et al. (2002) may overestimate the depleted mantle Hf isotope values, especially over the first 2.0 Ga of the Earth history. The curve also gives a high ε_Hf value of c. 1.3 at 4.56 Ga, using the CHUR constraints of Bouvier et al. (2008).

(2) The initial mantle depletion took place before c. 3.8 Ga ago, and increasing and significant depletion is recorded in zircons before 3.0 Ga. Therefore, all the curves that treat depleted mantle and continental crust as complementary reservoirs with the crust growing through time since a major growth episode around 3.0 Ga (e.g. Taylor & McLennan 1995; Collerson & Kamber 1999) may underestimate the depleted mantle values before 3.0 Ga.

Pietranik et al. (2009) used intersections between trends for the evolution of unmodified crustal sources (see the next section) and peaks of ages of crust formation to obtain Hf isotope ratios of the depleted mantle. The ages of new crust formation were constrained by vertical trends on U–Pb and ε_Hf plots, which are often characterized by stepped increases in the maximum values of ε_Hf. This approach provides values of the mantle from which most of the continental crust was derived and it overcomes the difficulty that most of the zircons yields contaminated ε_Hf values. Unexpectedly perhaps, the reconstructed evolution of the depleted mantle is linear, it has present-day ¹⁷⁶Lu/¹⁷⁷Hf = 0.0393 and ε_Hf of c. +15.9, and it is best approximated by the function ε_Hf = (−3.85 ± 0.1)t + (15.9 ± 0.2) (R² = 0.99), where t is the depleted mantle age in billion years. It intersects chondrite evolution (CHUR) at c. 4.0 Ga, consistent with point (2) above. The linearity implies that the Hf isotope composition of the depleted mantle was subsequently little affected by the generation of younger continental crust. Either the rates of crust generation and recycling were broadly balanced (Armstrong 1981), or the crust was generated from a relatively large volume of geochemically homogeneous, and previously depleted mantle (Tolstikhin et al. 2006).

The low Lu/Hf material responsible for the early depletion of the mantle could have been crust that stabilized at c. 4.0 Ga, or an early enriched reservoir (EER) that resulted in an additional depletion episode before the present continental crust was formed (c. 4.4 Ga; Boyet & Carlson 2006; Tolstikhin et al. 2006; Shirey et al. 2008). In the former, crust generation was responsible for the degree of depletion observed, and so the volume of mantle involved must have been relatively small. For the latter, the volume of mantle could have been much larger (up to whole mantle, Tolstikhin et al. 2006), but the EER must have been sequestered away from well-mixed portions of mantle. Figure 8 shows a good agreement between the mantle curve based on zircon data and that representing a two-stage model of mantle depletion (EER extraction followed by continental crust extraction). In summary, the linear evolution of the depleted mantle appears to require early extraction of an enriched reservoir that leaves high Lu/Hf depleted mantle that is in turn little changed by the subsequent extraction of the continental crust. If correct, there is no straightforward link between the isotope evolution of the depleted mantle and the rates of generation of the continental crust.

Composition of the oldest crust

The composition of the early crust remains controversial. The presence of Hadean zircons suggests that granitic magmas, and thus felsic crust, were present at that time, but the key question is the composition of the crust and mantle from which such granitic magmas were derived. Lu/Hf ratios decrease with increasing differentiation and they therefore provide a tool with which to characterize variably differentiated crustal reservoirs (Fig. 4). On a plot of ¹⁷⁶Hf/¹⁷⁷Hf v. crystallization age (Fig. 3) a single crustal reservoir evolves along a straight line, and the slope of that line corresponds to its Lu/Hf ratio, and hence to its bulk composition.

The critical step is to recognize arrays representing single sources that have not been affected by mixing with older or younger crustal components, as was also required for evaluating the Hf isotope ratios of the depleted mantle (Fig. 6). One approach is to screen out zircons from magmas that contain contributions from sedimentary source rocks, as they will tend to yield hybrid Hf isotope ratios. Accepting only zircons with mantle-like δ¹⁸O (5.3 ± 0.3‰) or close to mantle-like values (usually zircons with δ¹⁸O up to 6.5‰ are accepted to accommodate the analytical precision of ion microprobe analysis; Cavosie et al. 2005) allows us to distinguish sources unmodified by contamination with supracrustal material. Figure 9 shows that zircons with mantle-like δ¹⁸O form three distinctive peaks in histograms of model ages for a range of Lu/Hf ratios: two peaks occur in the Archaean (3.5 and 3.1 Ga) and one, broader peak occurs at c. 1.5–2.0 Ga. The zircons plot as three arrays on plots of U–Pb age v. initial ε_Hf with surprisingly consistent slopes corresponding to ¹⁷⁶Lu/¹⁷⁷Hf = 0.019–0.024 (i.e. Lu/Hf = 0.135–0.17), typical of a basaltic source. The two Archaean sources appear to coexist for over 0.5 Ga but then after c. 2.7 Ga the signal of the older crust seems to diminish and the younger Archaean source signal disappears after the 2.0 Ga source is formed. Surprisingly, only three general sources have been identified using zircons with mantle-like δ¹⁸O, despite including zircons from a number of areas (Australia, Slave craton, Mississippi River and Ural) and there is a lack of a clear 2.7 Ga old source despite extensive magmatism at that time. This may in part reflect the present shortage of O isotope analyses in the literature, and/or the selection of the curve for the evolution of the depleted mantle.

Fig. 9.

(a) Evolution of ε_Hf through time in zircons with mantle-like δ¹⁸O. The black line is the depleted mantle evolution based on zircons, as in Figure 8, and the grey line is the Griffin et al. (2002) evolution curve. The zircons form three distinct crustal arrays with slopes corresponding to ¹⁷⁶Lu/¹⁷⁷Hf = 0.019–0.024. The zircons included in (a) therefore also form peaks in the Hf model age histograms (b–d), consistent with many zircons coming from a single source. The model ages in (b–d) were calculated using the zircon-based mantle evolution curve; using the Griffin et al. (2002) curve would result in model ages higher by 100–200 Ma, but with a similar distribution of peaks. For the model age calculations the selection of ¹⁷⁶Lu/¹⁷⁷Hf in the source determines the slope of the arrays in (a). We included zircons that formed peaks in model ages for a range of ¹⁷⁶Lu/¹⁷⁷Hf ratios typical for mafic to intermediate magma sources, and including zircons with model ages calculated using lower or higher ¹⁷⁶Lu/¹⁷⁷Hf had minimal effect on the slopes of the crustal arrays. The zircon data are from Kemp et al. (2006), Pietranik et al. (2008) and Wang et al. (2009).

Another way to identify evolution arrays of single crustal sources is to select areas where one source was reworked for over c. 0.5 Ga (the time needed to obtain a reasonable regression) without the addition of new crust, and to determine the slope of the evolution array formed by zircons with the highest ε_Hf. Figure 10a shows the evolution array constrained for South American zircons (Willner et al. 2008), and although this approach remains difficult to verify statistically, the slope of the array is again consistent with mafic sources for these magmas. The dominance of implied mafic sources for ancient zircon arrays is consistent with the expectation that new crust generated from mantle peridotite is basaltic, rather than being similar to average present-day andesitic crust.

Fig. 10.

Crustal arrays defined in zircon age v. ε_Hf diagrams highlighting zircons that crystallized from magmas extracted from the youngest (a) and the oldest (b) crustal sources sampled by South American zircons (Willner et al. 2008).

Similar approaches can be applied to the oldest zircons known, from Jack Hills in Western Australia. The solution Hf isotope data of Amelin et al. (1999) define an array whose slope corresponds to the evolution of a source with ¹⁷⁶Lu/¹⁷⁷Hf c. 0.022, which is most typical of a mafic crustal composition (Amelin et al. 1999). Other studies have identified zircons with lower apparent ε_Hf values that might imply the remelting of felsic, tonalite–trondhjemite–granodiorite-like (TTG-like) sources with ¹⁷⁶Lu/¹⁷⁷Hf values <0.01 (Harrison et al. 2005, 2008; Blichert-Toft & Albarède 2008). However, for samples younger than 4.0 Ga there is little evidence that such source regions persisted, which is unexpected given the dominance of TTG complexes in Archaean terrains. The Archaean zircons with low ε_Hf appear to plot on ‘mafic' arrays with slopes corresponding to ¹⁷⁶Lu/¹⁷⁷Hf of 0.022 (Fig. 10b). The evidence from zircons is that both mafic and felsic crust were present in the Hadean and the early Archaean, and that the mafic crust may have been predominant. This is consistent with heat production arguments that felsic crust is unlikely to have been thicker than c. 10 km, whereas mafic crust might have been c. 40 km thick (Kamber et al. 2005).

Peaks of ages, preservation and crust generation

Figure 1 illustrates the peaks of the ages of rocks that, on the basis of their mantle-like initial isotope ratios, are thought to represent new continental crust (McCulloch & Bennett 1994; Condie 1998, 2005). As indicated above, there is considerable debate over the extent to which these peaks are representative of the evolution of the continental crust or are a consequence of preferential preservation in the geological record. If they are assumed to be representative, the presence of alternating peaks and troughs of ages implies dramatic changes in the rates of crustal production from periods in which unusually large volumes of new crust were generated to periods of relatively little new crustal production. It is difficult to envisage how the global rate of crust generation would vary markedly if the generation of new continental crust is in response to plate tectonics. The apparent peaks of crust generation have therefore been attributed to deep-seated thermal anomalies in the mantle, and the emplacement of superplumes (Stein & Hofmann 1994; Albarède 1998) or the triggering of mantle avalanches (Condie 1998; Nelson 1998). One difficulty is that intra-plate magmatic rocks are not commonly recognized in Archaean granite–greenstone terranes. Another is that experimental evidence for the generation of the TTG indicates that they are likely to have been generated from altered basalt (Foley et al. 2003; Brown & Rushmer 2006; Clemens et al. 2006), and it is less clear how hydrothermally altered basalt can be taken down to the site of partial melting in an intra-plate setting.

Campbell & Allen (2008) summarized the crystallization ages from c. 7 000 zircons, most of which were sampled from recent sediments from over 40 rivers worldwide (Fig. 11). They too yield marked peaks in crystallization ages and, again assuming that they are representative of igneous events in the sedimentary provenance, they would reflect periods of increased magmatic activity in the evolution of the continental crust. Some of the peaks of crystallization ages, most notably at 1.9 and 2.7 Ga, are coincident with the periods in which relatively large volumes of new crust were generated (as indicated in Fig. 1). However, this is much less marked in the last 1.5 Ga, when the peaks of crystallization ages appear to represent pulses of magmatic activity that are not associated with the generation of significant volumes of new crust. This is consistent with models in which the volumes of new continental crust generated decrease exponentially with time as the Earth cools (e.g. Walzer & Hendel 1997; Condie 2000; Grigné & Labrosse 2001), with less and less new crust being generated in the younger magmatic events.

Fig. 11.

Plots of the age distribution of relative volumes of juvenile continental crust (from Condie 2005, and as in Fig. 1), and of crystallization ages for over 7000 detrital zircons (Campbell & Allen 2008). The peaks in the zircon crystallization ages are similar to the ages of supercontinents. The crust generation rate curve illustrates a model in which the volume of new crust generated decreases with decreasing age, and the lightly shaded peaks schematically illustrate the relative volumes of new crust that might consequently be associated with each peak of zircon crystallization ages.

A number of studies have pointed out that the peaks of crystallization ages occur at times when there were supercontinents on the Earth's surface (Taylor & McLennan 1995; Condie 1998; and see Campbell & Allen 2008; Rino et al. 2008, for summaries). It is less clear that these should have been periods of unusual volumes of magmatism, but it may be that rocks enveloped within supercontinents have more chance of being preserved. Brown (2007) categorized high-grade orogenic belts into high-, intermediate- and low-pressure high-temperature belts. He noted that the high-pressure belts were restricted to the last 600 Ma, and concluded that they reflect cold subduction as observed at present along convergent margins. Intermediate- to low-pressure high-temperature rocks are preserved dating back to the late Archaean, and Kemp et al. (2007b) pointed out that their ages are grouped in clusters similar to the peaks of crust generation illustrated in Figure 1. The implication is that periods of granulite-facies metamorphism are in some way linked to the processes of crust generation as suggested by Kemp et al. (2007b), and/or the peaks of the ages of crust generation and granulite metamorphism are themselves a function of the unevenness of the continental record.

One question is the extent to which the observed peaks of ages are a robust record of the major magmatic events in a particular area. This can be assessed by comparing the peaks of ages sampled by zircons in sediments of different ages. Analyses are available on over 11 000 detrital zircons from Australia, from sediments that range in age from Recent to Archaean (Fig. 12). The age distribution for all the zircons considered together is marked by a sharp peak at c. 2.7 Ga, and smaller peaks at c. 1.6–1.8, c. 1.2 and 0.5 Ga. These four peaks are more clearly seen in the zircons from sediments deposited in the last 50 Ma, but then the zircons in sediments through much of the Phanerozoic are dominated by the peaks at c. 1.2 and 0.5 Ga. The age peaks at 1.6–1.8 Ga begin to appear in zircons from the Ordovician and Cambrian sediments, and then the older peaks dominate in the older sediments. It appears that whereas different age peaks dominate in zircons from sediments of different ages, all of the four main peaks are sampled by the youngest sediments, and there is little evidence of age peaks that were sampled only by older sediments and then not sampled subsequently. The implication is that the age peaks sampled in the youngest sediments are a robust record of the major magmatic events recorded in Australian geology. The question is then the extent to which those age peaks are representative of the events that shaped the geology of this continent.

Fig. 12.

A summary of the distribution of crystallization ages from detrital zircons from sediments of different ages in Australia. For the most part the peaks of ages older than the host sediments do not change markedly in sediments of different ages, which implies that the age distributions obtained provide a robust representation of the preserved geological record of Australia. The sketch maps illustrate the locations of the dominant units of different ages.

Hawkesworth et al. (2009) explored ways in which the peaks of ages might be artefacts of preservation in the geological record, and preferential preservation might be linked to the development of supercontinents. The fossil record is biased by the unevenness of the geographical and stratigraphical sampling effort and inequality in the rock record available for sampling (Smith 2007). It seems increasingly likely that a similar unevenness biases the record of the generation and the evolution of the continental crust.

Supercontinents are regarded as an inevitable consequence of plate tectonics (e.g. Dalziel 1992); they are the outcome of global tectonics and they come together in compressive phases associated with subduction. The igneous record associated with the development and break-up of supercontinents is therefore one of subduction-related magmatism, collisional orogens and crustal melting, and subsequently extensional magmatism (Fig. 13). At issue is the volume of magma generated in each phase, and the extent to which the magmatic record of each phase will be preserved and specifically will be represented in the record of detrital and inherited zircons.

Fig. 13.

A schematic cross-section of convergent, collisional and extensional plate boundaries associated with supercontinent cycle showing estimated amounts (in km³ a) of continental addition (numbers in parentheses above Earth surface) and removal (numbers in brackets below surface). Data from Scholl & von Huene (2007, 2009). The volume of continental crust added through time via juvenile magma addition appears to be compensated by the return of continental and island arc crust to the mantle. Scholl & von Huene (2007, 2009) estimated that the long-term global average rate of arc magma additions is 2.8–3.0 km³ a (see also Franz et al. 2006). The total volume of crustal material moved into the mantle at convergent and collisional boundaries is around 3.2 km³ a. This rate is sufficient that if plate tectonics has been operating since around 3.0 Ga (see Cawood et al. 2006) then a volume equal to the total current volume of continental crust would have been recycled into the mantle (Scholl & von Huene 2007). The implication is that the net growth of continental crust at the present day is effectively nil, and convergent plate margins are sites of crustal recycling and reworking, as well as continental addition.

The composition of the continental crust appears to be dominated by geochemical signatures associated with compressive plate margin magmatism (Taylor 1967; Rudnick 1995; Hawkesworth & Kemp 2006a,b), yet rocks from this setting have relatively poor preservation potential in the geological record. The data compiled by Scholl & von Huene (2007, 2009) highlight that the global rates of removal of continental and island arc crust through subduction into the mantle are similar to the rates at which crust is generated at modern magmatic arcs (c. 2.5 km³ a⁻¹, Fig. 13). Clift & Vannucchi (2004) and Clift et al. (2009) used different datasets to reach similar but slightly higher values for the generation and recycling of continental material. Mineral deposits that form predominantly in convergent margin settings, such as epithermal and porphyry copper deposits and orogenic gold deposits, are generally less than 100 Ma old, as a result of rapid exhumation and erosion (Kesler & Wilkinson 2006; Bierlein et al. 2009), and this too is consistent with the poor preservation potential of arc magmatic rocks in the geological record.

Collisional magmatism, by contrast, is dominated by partial melting of the pre-existing crust. It is granitic, and although the volume generated may be small relative to some other tectonic settings (Fig. 14), it will tend to be preferentially protected within the enveloping supercontinent and it will have good preservation potential in the geological record. Denudation of the orogenic belt formed within the core of the supercontinent as a result of craton assembly will be a major source of detritus throughout the supercontinent and along its margins (e.g. the late Mesoproterozoic ‘Grenville' and late Neoproterozoic ‘Pan-African' age detritus that is recorded in Rodinia and Gondwana; Cawood et al. 2007). In contrast, the extensional phase is dominated by mafic magmatism; for example, the flood basalts associated with the break-up of Gondwana (Storey 1995; Hawkesworth et al. 1999). This phase is unlikely to result in large volumes of zircons, and the rocks may in any case be relatively sensitive to erosion into the oceans.

Fig. 14.

The volumes of magma generated (continuous line), and their likely preservation potential (dashed lines), may vary in the three stages associated with the convergence, assembly and breakup of a supercontinent (after Hawkesworth et al. 2009). The preservation potential in the first stage is greater at margins where the subduction zone retreats oceanward to form extensional basins than at margins where the subduction zone advances toward the continent. Thus, peaks in the crystallization ages that are preserved (shaded area) reflect the balance between the magma volumes generated in the three stages and their preservation potential.

It is argued that the record of magmatic ages is likely to be dominated by periods when supercontinents assembled, not because this is a major phase of crust generation but because it provides a setting for the selective preservation of crust. The preservation potential, particularly for crystallization ages of zircons, is greater for late-stage collisional events as the supercontinents come together, rather than for subduction- and extension-related magmatism (Fig. 14). Such an interpretation is consistent with the data for the last 1.5 Ga (Fig. 1), but at 1.9 and 2.7 Ga the peaks of ages of crystallization also match up with striking peaks of crust generation. The processes of crust formation dominate the geological record at these times. The events at 1.9 Ga are still much debated, but by 2.7 Ga the processes that shaped the geological record were strikingly different.

The end of the Archaean (2.9–2.6 Ga) is the time of the stabilization of the Archaean cratons as preserved at present. This reflects a particular stage in the cooling of the Earth (e.g. Vlaar et al. 1994; King 2005; Korenaga 2006), and the predictable consequence is that relatively large volumes of rock are preserved from that time (Fig. 1). One possibility is that the 2.9–2.6 Ga rocks are representative of the rocks and associations that formed earlier, although for how far back in time is not well constrained, but is simply trapped and preserved by the accident of history at this stage in the cooling of the outer part of the Earth. For rocks formed earlier in the Archaean, the preservation potential is extremely poor (as a result of increased recycling rates in a hotter Earth), and we infer that the record is therefore much less likely to be biased by the tectonic setting in which the rocks were formed.

Thus it is envisaged that the late Archaean marks the transition from a period of relatively poor preservation to one in which the geological record is biased by the tectonic setting in which the rocks were formed. It follows that for events older that 2.9 Ga zircons may provide the best available record, because rocks from different settings then have similar preservation potential, albeit extremely poor, and the record is less biased by the controls on preservation that mark post-Archaean magmatic processes.

An important inference that can be drawn from the compilations of volumes of continental generation and recycling (Fig. 13) is that the current volume of continental crust would have been recycled back into the mantle over the last 2–3 Ga and thus this volume of continental crust must have existed prior to this time.

The igneous record: insights from granites

Igneous rocks are involved in the generation of new continental crust, and recent isotope studies have focused on (1) the relative contributions of new and pre-existing crust in the generation of the granites, and on tracking the geodynamics of crustal evolution, and (2) unmixing granite sources to pinpoint times of crustal growth. They are well illustrated using the Tasmanides of southeastern Australia (Kemp et al. 2006, 2009b), a 700 km wide segment of an 18 000 km long orogenic system that developed along the palaeo-Pacific margin of Gondwana from the Neoproterozoic to early Mesozoic (Cawood 2005).

Within the Tasmanides, the Lachlan Fold Belt has a prominent place in the development of ideas about the petrogenesis of granite. It is where the influential classification scheme of I- and S-type granites was developed (Chappell & White 1974, 1992), and it is an area that has been at the centre of debates about the balance of new and pre-existing crust involved in the generation of granite, and how those are best determined (Gray 1984; Keay et al. 1997; Collins 1998, 1999; Chappell et al. 1999). This is clearly a central issue for discussions of when and how the continental crust was generated. The I- and S- classification scheme was developed at a time when granites were thought to be windows on their source regions, and those source regions could be inferred to be infracrustal (i.e. broadly igneous) or supracrustal (typically sedimentary) material. Granites were regarded as a way of looking into compositional variations deep in the crust. One difficulty was over the introduction of heat to induce partial melting. In many models this would come from the mantle and therefore include contributions from mantle-derived magmas. Furthermore, the Lachlan Fold Belt granites define an apparently simple, overlapping array on the ε_Nd v. initial ⁸⁷Sr/⁸⁶Sr diagram (McCulloch & Chappell 1982), which has been almost universally used to infer a large-scale mixing process between primitive magmas from the depleted upper mantle and evolved crustal end-members (e.g. Gray 1984; Faure 1986). The implication is that single granites were unlikely to have been derived from single source regions.

The Lachlan Fold Belt has two main components, a monotonous succession of mature (quartz- and clay-rich) Ordovician turbidites and a large volume of granitic rocks. The turbidites apparently accumulated on an oceanic substrate and were subject to episodic deformation, low-grade regional metamorphism and massive igneous intrusion from c. 450 to 340 Ma (Gray & Foster 2004). Such granite–turbidite belts appear strikingly different from orogenic tracts developed in collisional belts through a Wilson cycle of ocean opening and closing, and they form part of an accretionary orogenic belt (Coney 1992; Cawood et al. 2009). The Lachlan Fold Belt is thought to have occupied a back-arc setting for much of its evolution, behind a broadly eastward migrating subduction zone (Collins 2002a,b; Gray & Foster 2004; Cawood 2005; Foster et al. 2005).

Kemp et al. (2007a) reported U–Pb, Hf and O isotope measurements on magmatic zircons from three granite suites in the Lachlan Fold Belt. The striking feature of the Hf isotope data is that zircons from single whole-rock samples exhibited a spectrum of ε_Hf values of up to 10 ε_Hf units. Such variations within a single sample are reconciled only by the operation of open-system processes that are capable of modifying the ¹⁷⁶Hf/¹⁷⁷Hf ratio of the magmas from which the zircons precipitated. The polarity of Hf isotope changes during zircon growth can be determined from the isotope changes during the growth of single crystals, and by pairing the isotope variations with trace element ratios (e.g. Th/U) that are proxies for the degree of differentiation. In most cases, ¹⁷⁶Hf/¹⁷⁷Hf decreases with increasing differentiation (see Fig. 4) as would be induced by addition of an unradiogenic (continental crust-like) component during crystallization. Significantly, however, the zircons from both I- and S-type granites have Hf isotope ratios that trend back towards values similar to those in mantle-derived magmas. The in situ analyses of Hf isotopes in zircons are most consistent with there being mantle contributions in the generation of both I- and S-type granites.

Figure 15 summarizes the variations in ε_Hf in zircons from granitic rocks with their age of crystallization across the Tasmanides (Kemp et al. 2009b). The changes from compressional to extensional tectonic regimes are marked by changes in zircon ε_Hf (and whole-rock ε_Nd, not shown). Thus, for example, crustal reworking after back-arc closure is registered by the marked decrease in zircon ε_Hf that follows major crustal thickening episodes in each Tasmanide terrane (Fig. 15). These inflections signify the emplacement of S-type rocks generated under granulite-facies conditions largely from metasedimentary precursors (White & Chappell 1977). Subsequent extension is accompanied by increases in ε_Hf and an accompanying change from peraluminous S-type to metaluminous I-type compositions, presumably in response to a waning sedimentary contribution, as suggested by a steady decrease in zircon δ¹⁸O (Kemp et al. 2009b). The links between the shape of the isotope–time trends and the pattern of compressional and extensional events in the Tasmanides (Fig. 15) highlight the feedback between tectonic activity and magma source. In detail, Hf isotope ratios of zircons have been used to estimate the mantle contributions in different granite bodies generated at different tectonic stages. These range from 30–40% in the S-types analysed to c. 70% in the I-types, and up to 90% in the A-types, and clearly these represent the volumes of new crust generated in the Lachlan Fold Belt orogenic episode (Kemp et al. 2009b). Unexpectedly, because the S-type granites are more voluminous than the other granite types, a large proportion of the new crust appears to have been generated during S-type granite magmatism.

Fig. 15.

The isotope and tectono-magmatic evolution of the Tasmanides defined by the ε_Hf values of granite-hosted zircons (after Kemp et al. 2009b). Filled diamonds represent analyses from mafic units. The grey shaded time slices correspond to major contractional episodes, as follows: HB, Hunter–Bowen; K, Kanimblan; T, Tabberabberan; Bo, Bowning; Be, Benambran (early phase); D, Delamerian. The timing of these events is taken from Landenberger et al. (1995), Collins & Hobbs (2001), Collins (2002a), Gray & Foster (2004), Foden et al. (2006) and Cawood & Buchan (2007).

The second aspect is in using inherited and detrital zircons to pinpoint periods of crustal growth. The detrital zircons in greywackes from the Lachlan Fold Belt and the inherited zircons in the granitic rocks yield similar distributions of crystallization ages (Kemp et al. 2006; Fig. 16). There is a marked peak at 500–650 Ma and another at 0.9–1.2 Ga that manifest Pan-African and Grenvillian phases of supercontinent-related orogenesis, respectively, and then there is a scattering of ages back to 3.5 Ga. The peaks of zircon crystallization ages are taken to reflect peaks of magmatic activity prior to the formation of the Lachlan Fold Belt, and the initial question is the extent to which these reflect periods of new crustal growth, or crustal reworking primarily through remelting of the pre-existing crust.

Fig. 16.

A histogram of ages obtained from inherited and detrital zircons from c. 430–380 Ma old granites and Ordovician turbidites in the Lachlan Fold Belt (Kemp et al. 2006). The crystallization ages have peaks at c. 500 and 1000 Ma, whereas the Hf model ages (see Fig. 3), here termed crust formation ages, are much older with peaks at c. 1.9 and 3.3 Ga. The peaks of model ages in zircons with δ¹⁸O < 6.5‰ indicate when new crust in the areas sampled by these zircons was generated. The model ages of zircons with δ¹⁸O > 6.5‰ contain a contribution from supracrustal rocks, and they are therefore more likely to represent hybrid model ages.

Hf model ages represent the times when the crustal source rocks for granitic magmas were themselves derived from the mantle (Fig. 3). Kemp et al. (2006) used O isotopes to identify zircons that crystallized from magmas that contained a contribution from sedimentary source rocks (those with δ¹⁸O > 6.5‰), as these are likely to have hybrid model ages that are unlikely to represent discrete crust-forming events. Zircons with δ¹⁸O < 6.5‰ are thought to have been derived from magmas derived from igneous source rocks, and hence to have model ages that are more likely to reflect periods of crust generation. The results are very striking. Zircons with δ¹⁸O < 6.5‰ define two marked peaks of Hf model ages at 1.9 and 3.3 Ga (Fig. 16, using the depleted mantle model of Griffin et al. 2002), and these are taken to be periods of generation of significant volumes of new continental crust in the source region. In contrast, the zircons with δ¹⁸O > 6.5‰ yield a range of model ages. Many are intermediate between 1.9 and 3.3 Ga, they have a broad peak at 2.1–2.2 Ga, and so they presumably represent mixtures of rocks generated at 1.9 and 3.3 Ga. The oxygen isotope data therefore allowed Kemp et al. (2006) to conclude that the 2.1–2.2 Ga age peak represented mixing in the sedimentary environment, rather than a real crust-forming event. More generally, with this approach we can now contrast the evolution of the igneous and sedimentary reservoirs in the continental crust, and compare the information from zircons with the models for the evolution of the continental crust based on Nd isotope ratios in shales (Allègre & Rousseau 1984).

The sedimentary record: erosion models and continental maturation

Fine-grained continental sediments sample the available crust at the time of deposition. They are widely used to obtain average abundances of the upper crust for insoluble elements, and so there is considerable interest in using Nd and Hf isotopes in sediments to constrain models for the evolution of the crust.

Allègre & Rousseau (1984) modelled Nd isotopes in shales of different ages in terms of the growth of the continental crust. On a plot of the model age of the sample against its sedimentation age, or crystallization age in the case of zircons, new crust that is then reworked in subsequent events results in horizontal arrays (Fig. 17). Samples that have younger model ages in younger sediments, as seen for Nd isotopes in shales, indicate that new crust was generated in younger events and it has then been sampled by the younger sediments. As the slope of the data array flattens out, it indicates that with time less and less new crust has been generated. Allègre & Rousseau (1984) assumed that on average new crust was generated every 500 Ma, and they evaluated the links between the sediments analysed and their source rocks using an erosion factor ‘K'. K relates the model age of the sediments analysed to the average model age of their source rocks, given that some source terrains are more susceptible to erosion than others. Models were evaluated for different values of K, largely because K has been difficult to constrain (Allègre & Rousseau 1984; Goldstein & Jacobsen 1988; Jacobsen 1988; Kramers & Tolstikhin 1997; Tolstikhin & Kramers 2008).

Fig. 17.

(a) Model ages v. crystallization ages of detrital zircons from five recent sediments from the Frankland River in SW Australia (Dhuime et al. 2009). The bold diamonds are the means of the zircon data (small dots) grouped into five main periods of zircon crystallization: 3.4–3.0, 2.8–2.5, 2.4–1.4, 1.35–1.0 and 0.7–0.4 Ga. The average values for zircons of different ages define a similar trend through time to the model Nd ages of shales of different ages from the Australian continent (dashed curve). The latter was used by Allègre & Rousseau (1984) to model the evolution of the continental crust using different values of the erosion constant ‘K' (see text). The inset illustrates that new crust would plot on the 1:1 line, and new crust subsequently defines a horizontal array retaining the same model age through younger geological events. Displacement down the diagram (i.e. to younger model ages) requires a contribution from younger crustal material. (b) The relative contribution of zircons from the Yilgarn block in the zircon populations in recent sediments plotted against distance along the Frankland River (grey envelope curve). These are compared with the proportion of Archaean source rock available in the catchment for each sediment sample (open squares). The calculated K values for each sediment, and the change in elevation down the river, are also shown. The inflexion in the river's profile at about 80 km from the coast is responsible for the increase of K. The high values of K close to the contact with the Yilgarn are subject to large errors, and are poorly understood. Values of K = 4–6 appear representative of mature river systems that sample large areas of continental crust.

It remains difficult to assess independently the relative contributions of sources of different ages in a sample of sediment. One way forward is to combine Hf isotopes in zircon with Nd isotopes in whole-rock samples. The distribution of Hf model ages in detrital zircons in a sediment offers insight into the proportions of different source terrains that have contributed to the bulk sample, and hence on the actual value of K. However, zircon is a heavy mineral and the Nd budget in sediments is dominated by the clay fraction (see Vervoort et al. 1999), and it has to be established how well these records compare.

Dhuime et al. (2009) undertook a Hf and Nd isotope study on samples of recent sediment from the Frankland River in southwestern Australia, building on the earlier study of Cawood et al. (2003). The Frankland River is c. 320 km in length and the catchment area is c. 4630 km². It drains two terranes with different zircon crystallization and model age distributions, the Archaean Yilgarn craton and the Proterozoic Albany–Fraser mobile belt. Thus it offers an opportunity to compare the distribution of the rocks sampled in different sediments and the proportions of those rocks in the catchment area for each sediment. The Hf and U–Pb age data on zircons from four recent sediments are summarized in Figure 17. Although there is a large range of Hf model ages in any group of zircons with similar U–Pb ages, it is striking that the trend of the average value of the Hf model ages through time is broadly similar to that for Nd model ages in Australian shales (Allègre & Rousseau 1984). This suggests that the distribution of zircon data may be used to constrain values of K for different sediment samples.

Dhuime et al. (2009) used the model age and the crystallization age data on detrital zircons to calculate the contribution from the Yilgarn craton and the Albany–Fraser belt in each sample, and to compare that with the proportion of those two units in the catchment again for each sample. Overall there is a downstream decrease in the proportion of Yilgarn material in the recent sediments (Fig. 17b, grey curve). Moreover, the K values increase with distance from the Yilgarn craton, and with the gradient of the river profile (K varies from c. 4–6 to c. 15–17, Fig. 17b). Samples with K c. 9–10 and c. 15–17 are from below the inflection that reflects a weak escarpment at c. 80 km from the coast associated with Miocene–Pliocene uplift (Cawood et al. 2003) (Fig. 17b). The steepening of the river's profile has resulted in preferential erosion of material from the Albany–Fraser belt, and hence an increase in the calculated K values. In turn the ‘stable' segment of the Frankland River is best sampled above the escarpment at distances of 100–150 km from the coast. The implication is that K values of 4–6 are representative of mature river systems that sample large areas of continental crust, and models using values of K = 2–3 result in biases towards younger ages in the global models of crust formation and evolution through time (Allègre & Rousseau 1984; Goldstein & Jacobsen 1988; Jacobsen 1988; Kramers & Tolstikhin 1997). It is encouraging that Nd and Hf isotopes can now be combined to evaluate the erosion constant K and how it varies with changes in discharge rates and relief.

Synthesis

New analytical approaches have offered new insights into the generation and evolution of the continental crust. This remains the key archive of how processes and conditions have changed during the evolution of the Earth. The oldest rock is 4 Ga old, and only 7% of the preserved continental crust is older than 2.5 Ga. Yet most of the crust is generally inferred to have been generated by that time (see Fig. 11), and the challenge is to decipher the major events in the generation of the crust from this small fragment of the geological record. From 4.4 to 4.0 Ga the only samples are tiny crystals of detrital zircons, and therefore much of the discussion has focused on the information available from the zircon crystals themselves and the silicate inclusions within them. Zircons are the basis for the geological time scale, in that they yield precise crystallization ages, and they have been increasingly used to determine crust formation ages using Hf isotopes. The continental record is marked by peaks in ages of crystallization, which in turn imply periods of enhanced magmatic activity, and by peaks in ages of crust formation, which might be taken to reflect periods of enhanced crust generation. However, an alternative interpretation is that these peaks of ages are not a true record of such igneous processes, but instead an artefact of preservation. The geological record is far from complete, and the fossil record is biased by the nature of the fossils, sedimentary facies and the nature of the rock record. The peaks of crystallization ages also mark the times of supercontinent formation, and it is increasingly understood that the preservation potentials of rocks generated in different tectonic settings are very different. It may simply be that the development of a supercontinent offers markedly improved preservation potential for magmas formed at such times (Hawkesworth et al. 2009).

The nature of the crustal source rocks of the magmas from which Hadean and early Archaean zircons crystallized can be inferred from thermal considerations (Kamber et al. 2005), and the estimated source Lu/Hf ratios. Both approaches indicate that the early crust was predominantly mafic in composition, but that there was a felsic component, as sampled by some of the magmas from which the Jack Hills zircons crystallized (Harrison et al. 2005, 2008; Blichert-Toft & Albarède 2008). The early crust is likely to have had a bimodal distribution in silica, and as such to have been different from geological younger subduction-related associations. For rocks formed in the Hadean and early Archaean the preservation potential was extremely poor, presumably because of meteorite bombardment and increased recycling rates in a hotter Earth. In many cases detrital and inherited minerals, such as zircon, remain the main geological archive, and there is little sense that this early geological record is biased by the tectonic setting in which the rocks were formed. The end of the Archaean (2.9–2.6 Ga) is the time of the stabilization of the Archaean cratons as preserved at present. One implication is that this reflects a particular stage in the cooling of the Earth, and that is why relatively large volumes of rock are preserved from that time (Fig. 1). It may be that the 2.9–2.6 Ga rocks are representative of the rocks and associations that formed earlier, although for how far back in time is not well constrained, but is simply trapped and preserved by the accident of history at this stage in the cooling of the outer part of the Earth.

Many models indicate that large volumes of continental crust were generated by the end of the Archaean, and that the volumes of new crust generated decreased markedly since that time (e.g. Fig. 11). However, these arguments were in part based on the evolution of the depleted mantle that was attributed to the generation of the continental crust. Yet if the depleted mantle is linear in its isotope evolution (e.g. Fig. 8), and the depletion is attributed to early differentiation events, then it cannot be used to monitor the development of continental crust. Instead, arguments in favour of large volumes of continental crust, and the likely thickness of felsic and mafic crust, before the end of the Archaean rely on thermal models for the decay in radiogenic heat production and the progressively cooling Earth (Davies 1999; Kamber et al. 2005). These are in turn consistent with recent estimates that the rates of crust generation and destruction along subduction zones are strikingly similar (Clift & Vannuchi 2004; Scholl & von Huene 2007, 2009). The implication is that the present volume of continental crust was established 2–3 Ga ago.

The Late Archaean marks a change, with the subsequent record controlled by the competing forces of preservation and generation (Fig. 14). Supercontinents were developed intermittently, and palaeomagnetic data indicate that blocks of crust moved laterally relative to one another (Cawood et al. 2006). Plate tectonics may have been active before 3 Ga, but it was dominant since the end of the Archaean. The geological record, however, is biased by peaks of ages linked to the development of supercontinents (Fig. 11), and, as we have argued elsewhere, by the development of granulite-facies rocks that are in turn difficult to destroy (Kemp et al. 2009b). One test of the links between the development of granulite-facies rocks and the stabilization of the continental crust is through Pb isotopes. Granulites are characterized by low U/Pb, and hence with time unradiogenic Pb isotope ratios (e.g. Rudnick & Goldstein 1990; Zartman 1990). There has been a lengthy discussion on the need to identify a low U/Pb reservoir to model the Pb isotope evolution of the crust and mantle (Rudnick & Goldstein 1990; Zartman 1990; Asmerom & Jacobsen 1993; Kramers & Tolstikhin 1997; Hofmann 2008). Some have invoked granulite-facies lower crust, and typical models suggest μ values (²³⁸U/²⁰⁴Pb) of 3–4 in the lower crust, 10–11 in the upper crust and 7–8 in the bulk crust. Thus, c. 30% of the continental crust may be granulite-facies rocks that are relatively difficult to destroy, and so help in the stabilization of the crust.

One implication of the different preservation potential of rocks generated in different settings is that different records should be preserved in sediments that survive from different settings. There is also increasing evidence that models for the evolution of the continental crust now need to integrate large-scale, plate-tectonic cycles that result in the development and destruction of supercontinents, and shorter-lived 25–100 Ma cycles that differ depending on the extent to which the plate margin is advancing or retreating. Such cycles have been identified in the western American Cordilleras, an advancing orogen where the overriding plate exerts the dominant control (DeCelles et al. 2009). In sharp contrast, the Tasmanides represent the retreating orogens of the western Pacific, which are controlled by the lower (subducting) plate and where new crust is generated, and stabilized, in back-arc settings (Kemp et al. 2009b). There is now considerable scope to establish the evolution of single orogens and to use isotopes in igneous rocks to develop more detailed models of the geodynamics of crust generation and evolution.

Acknowledgments

We thank R. Strachan for inviting this review, and his subsequent patience during the writing of the manuscript, and M. Whitehouse, S. Daly and A. Prave for their thorough and supportive reviews. C.H. gratefully acknowledges support from the NERC (NE/E005225/1) and a Royal Society Wolfson Award, as does C.S. for an NERC Fellowship NE/D008891/1, and A.K. for an Australian Research Council Fellowship (DP0773029).

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The generation and evolution of the continental crust

Abstract:

Old problems and new ways forward

Terms

The timing of events: crystallization and crust formation ages

Variations in Sm/Nd and Lu/Hf ratios and the composition of initial continental crust

Zircons as archives of the continental crust

Differentiation of the infant silicate Earth: formation of early depleted and enriched reservoirs

Composition of the oldest crust

Peaks of ages, preservation and crust generation

The igneous record: insights from granites

The sedimentary record: erosion models and continental maturation

Synthesis

Acknowledgments

References

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The generation and evolution of the continental crust

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