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LITHOS ELSEVIER
Lithos48 1999 vii
Contents
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LITHOS ELSEVIER
Lithos 48 1999 ix-xii
Preface
Composition, Deep Structure and Evolution of Continents
In the spring of 1997, the Department of Earth and Planetary Sciences of Harvard University and the Department of Earth, Atmospheric, and Planetary Science of the Massachusetts Institute of Technology jointly organized a series of lectures, seminars, and student assignments on a range of issues pertaining to the deep 'roots' of continents, which culminated in the International Symposium on Deep Structure, Composition, and Evolution of Continents, Harvard University, Cambridge, Massachusetts, Oct 15-17, 1997. The ensemble of manuscripts presented in this special volume captures the stimulating cross-disciplinary dialogue of this meeting, and we hope that it will both provide an update on recent research developments and serve as a starting point for research of the many outstanding issues. After its formation at mid-oceanic spreading centers, oceanic lithosphere cools, thickens, and subsides, until it subducts into the deep mantle beneath convergent margins. As a result of this continuous recycling process oceanic lithosphere is typically less than 200 million years old the global average is about 80 Myr.. In contrast, some parts of continental lithosphere still exposed at Earth's surface were formed more than 4 billion years ago the average age of continental rocks is about 2 billion years, and secular cooling can only explain part of the observations pertinent to continental structure. Indeed, ancient continents appear to be complex assemblages whose structure, composition and evolution can only be constrained by combining the results of geophysical imaging, geological mapping, petrological and geochemical analyses, and numerical modeling. A comprehensive, multi-disciplinary study of continents involves a wide range of length scales: tiny
PlI: S0024-4937 99 00049-3
rock samples and diamond inclusions may yield isotope and trace element signatures diagnostic for the formation age and evolution of parts of cratons, while geophysical techniques e.g., seismic and electromagnetic imaging, constrain variations of elastic and conductive properties over length scales ranging from several to many thousand kilometers. Integrating and reconciling this information is far from trivial and, as several papers in this volume document, the relationships between, for instance, formation age and tectonic behavior on the one hand and the seismic signature, heat flow, and petrology on the other may not be uniform but may vary both within as well as between cratons. These observations complicate attempts to determine the variations of one particular observable e.g., heat flow, lithosphere thickness, as a function of another e.g., crustal age. on the basis of global data compilations and tectonic regionalizations. Seismic imaging continues to provide unique constraints on the deep structure of continents. The wavelength of structure that can be resolved is constrained by the frequency of the seismic waves and the way they sample the study object. Seismologists therefore chose data sets and imaging techniques with a specific application in mind. In the first paper, Bostock demonstrates that high-frequency data can constrain mantle stratification and intra-crustal structures in spectacular detail, in particular if the analysis is not restricted to transmitted seismic waves but also involves waves that are converted and refracted at structure boundaries. The seismic data reveal that some structures in the shallow mantle are continuous to known shear zones in the crust, which has tremendous implications for our understanding of crest-
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Preface
mantle interaction during continental formation and deformation. The examples shown illustrate the recent successes in bridging the resolution gap between conventional 'passive source' imaging of mantle structures and the 'active source' refraction and reflection studies of the upper crust. Targeting much longer wavelengths, Simons and co-workers use surface waves traversing the Australian continent to demonstrate a strong relationship between the age of the crust and the seismic signature of the lithospheric mantle to ; 175 km depth. However, at larger depth and on lengthscales less than 1000 km the seismic signature varies significantly not only between geological provinces of different age but also within them. When using the seismic wavespeeds as a measure, the Archean cratons in Australia are probably less then 200 km thick. Within the uncertainties of the models, this agrees with new interpretations of seismic data for the Kaapvaal craton, as discussed by Priestley, but not with previous global models which have suggested that the roots beneath ancient cratons extend to depths in excess of 400 km. Important conclusions of the work presented here are that 1 continental deformation, for instance shortening, is not restricted to the crust but also involves the lithospheric mantle; 2 .the high wavespeed part of continental lithospheric mantle is probably thinner than inferred previously from vertically travelling body waves or form global surface-wave models; and 3 the seismic signature of ancient continents is more complex than expected from a uniform relationship with crustal age. Electromagnetic methods supplement seismic imaging in that it constrains different aspects of continental structure and composition. Whereas seismic imaging constrains bulk properties, electrical conductivity is primarily controlled by the presence in the rock matrix of constituents such as graphite, partial melt, and volatiles. It is particularly sensitive to the asthenosphere and has been used successfully to define the base of the resistive .cratonic lithosphere. Jones reviews the techniques used, along with their specific problems and promises, and shows that where both measurements exist, such as North America and the Baltic shield, the estimates of lithosphere thickness from magnetotelluric methods are often in excellent agreement with inferences from seismic imaging. Magnetotelluric data put constraints
on the relation between the solidus and the geotherm, and they can be used, if possible in concert with seismic data, to test hypotheses regarding the creation and subsequent development of continental lithospheric mantle. Heat flow potentially gives us direct insight into the thermal state of the continental lithospheric mantle. Many investigators have argued for a uniform inverse .relationship between the age of the crust and heat flow, which is often interpreted as evidence that the thickness of the conducting lithosphere increases with increasing age of the overlying crust. Drawing examples from the African and North American continents, Nyblade reviews such relationships. However, it is difficult to make accurate measurements, and the assumption of steady state on which many interpretations are based is not always justified because the characteristic time for heat diffusion is large compared to that of geological processes so that the latest thermal events may simply not yet be 'felt' at the surface. For thick conductive layers, in particular, such transient effects are hard to avoid. Moreover, one must account for the heat production within the crust, laupart and Mareschal conclude that the variations in surface heat flow between different geological provinces may not necessarily reflect differences in deep structure or the thickness of the lithosphere but that they can largely be explained by variations in heat production in the crust. If corrected for such effects, the heat flow appears to be rather constant across domains of different geological history. These transient effects and local concentrations of heat producing elements, which produce variations of observed heat flow within geological provinces of the same age, make it difficult to generalize on the basis of crustal age alone. The formation and subsequent evolution of continents is constrained by geological data and can be investigated by computer simulations. Shapiro and co-workers investigate under what circumstances the continental tectosphere can survive. In their first paper they discuss that compositional buoyancy is required to stabilize a tectosphere during initial formarion, but upon secular cooling the temperature dependence of the rheology becomes the dominating factor; if the viscosity is sufficiently high, compositional buoyancy is no longer required for stabiliza-
Preface
tion of the cold continental mantle. This concurs with De Smet and co-workers who model the formation and subsequent evolution of continental mantle and argue that the initial stage of melt extraction and tectosphere formation can be relatively short 50-100 million years, and that continued cooling prevents further differentiation after about 1 billion years. In their second paper, Shapiro and co-workers use constraints on seismic wavespeed to quantify the density structure of the tectosphere. While apparently not needed for maintaining the tectosphere's stability after its formation, the presence of a less-dense refractory mantle is required to explain why ancient continents contribute so little to the long wavelength gravity field. The gravity and seismic data are consistent with an isopycnic equal-density .structure enforced by compositional variations, but the effect of composition can be slightly less than envisaged in the original tectosphere model. However, if it is much less than the negative .thermal buoyancy the fit to gravity data degrades substantially. Geochronological studies provide significant insight into the age profile of lithospheric mantle and thus yield important constraints on processes of its formation and evolution. Both the Sm-Nd and Re-Os isotope systems reveal that ancient crust is often underlain by similarly aged lithospheric mantle. However, Pearson's review of the available data on the age of the continental lithospheric mantle illustrates the specific problems with each of the dating techniques and shows that, at present, there is no simple vertical age distribution for xenoliths from the Kaapvaal, Siberia, and Wyoming cratons. Diverse age distributions exist, creating complexities in reconstructing the deep lithosphere. Petrological and geochemical studies of samples of the deep lithosphere continue to yield new insights into the diversity of processes involved in the growth and evolution of continental roots. Often regarded as the paradigm for a cratonic root, the Kaapvaal craton in South Africa is characterized as old Archean ,. cold, and strongly depleted, perhaps by the extraction of komatiites. In contrast to sampies of post-Archean lithosphere, mantle xenoliths from Kaapvaal generally have higher orthopyroxene contents, a feature that is often attributed to melt depletion processes. However, this interpretation is
xi
debated and petrological analyses begin to reveal significant differences between cratons. Schmidberger and Francis bring to light the nature and diversity of mantle xenoliths from beneath the Canadian Arctic craton. They reveal that, like other Archean cratons, this region is underlain by a very refractory, melt-depleted root, which does not, however, possess an orthopyroxene-rich character. This may reflect a different style of ancient lithospheric growth, which adds to the pluriformity of present-day seismic and thermal structures of cratons as revealed by seismic imaging and interpretation of heat flow data. In addition to being less depleted, the postArchean lithospheric mantle appears to be susceptible to thinning and erosion by later tectonic and magmatic processes and may thus be less permanent than commonly assumed. In this volume, two papers report on such processes occurring today beneath younger portions of the continents. Stern and coworkers, in a study of South American mantle xenoliths, and Furman and Graham, in their study of East African lavas derived from the lithosphere, provide compelling evidence for the loss of continental roots during recent heating and melting events. Similar processes may have destroyed deep roots beneath the ancient Australian cratons, as discussed by Simons and co-workers in chapter two. Another consequence of re-heating and melting of the base of the lithosphere is the injection of melts derived from the lithosphere and asthenosphere into shallower parts of the peridotitic root. Glaser and co-workers report on melt infiltration in xenoliths from the post-Archean lithospheric mantle of Siberia, and Mukasa and Shervais examine products and processes of dike injection in the Balmuccia massif in Italy. Our picture of the deep lithosphere is not complete unless we also consider some of its non-peridotitic lithologies. Eclogites and pyroxenites are typically a minor component of xenolith suites, but they can reveal crucial insights into the tectonic and magmatic setting of craton formation. In this regard, Jacob and Foley's contribution focuses on Archean, diamond-bearing, eclogite xenoliths from Siberia and reveals that these samples are ancient fragments of island arc basalts that were once at the Earth's surface and have been subsequently recycled back into the mantle and stored in the cratonic root.
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Preface
Finally we would like to express our gratitude to the participants of the MIT-Harvard course for stimulating discussions and for the enthusiasm that resuited in the workshop at Harvard University and, eventually, in this volume. Richard O'Connell, in particular, is gratefully acknowledged for helping to organize the meeting and for his advice during the initial stages of the production of this volume. We also thank the many reviewers of the manuscripts presented here and the International Association of Seismology and Physics of the Earth's Interior, the International Association of Volcanology and Chem-
istry of the Earth's Interior, and the U.S. National Science Foundation for their support. Rob van der Hilst Massachusetts Institute of Technology Cambridge, MA 02139 USA Bill McDonough Har"ard Uni"ersi~. Cambridge, MA 02138 USA
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Lithos 48 (1999) 1-16
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Seismic imaging of lithospheric discontinuities and continental evolution M.G. Bostock Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, British Columbia, Canada V6T 1Z1
Received 12 March 1998; received in revised form 16 September 1998; accepted 21 September 1998
Abstract Discontinuities in physical properties within the continental lithosphere reflect a range of processes that have contributed to craton stabilization and evolution. A survey of recent seismological studies concerning lithospheric discontinuities is made in an attempt to document their essential characteristics. Results from long-period seismology are inconsistent with the presence of continuous, laterally invariant, isotropic boundaries within the upper mantle at the global scale. At regional scales, two well-defined interfaces termed H ( ~ 60 km depth) and L ( ~ 200 km depth) of continental affinity are identified, with the latter boundary generally exhibiting an anisotropic character. Long-range refraction profiles are frequently characterized by subcontinental mantle that exhibits a complex stratification within the top 200 km. The shallow layering of this package can behave as an imperfect waveguide giving rise to the so-called teleseismic Pn phase, while the L-discontinuity may define its lower base as the culmination of a low velocity zone. High-resolution, seismic reflection profiling provides sufficient detail in a number of cases to document the merging of mantle interfaces into lower continental crust below former collisional sutures and magmatic arcs, thus unambiguously identifying some lithospheric discontinuities with thrust faults and subducted oceanic lithosphere. Collectively, these and other seismic observations point to a continental lithosphere whose internal structure is dominated by a laterally variable, subhorizontal layering. This stratigraphy appears to be more pronounced at shallower lithospheric levels, includes dense, anisotropic layers of order 10 km in thickness, and exhibits horizontal correlation lengths comparable to the lateral dimensions of overlying crustal blocks. A model of craton evolution which relies on shallow subduction as a principal agent of craton stabilization is shown to be broadly compatible with these characteristics. 9 1999 Elsevier Science B.V. All rights reserved. Keywords: Seismic imaging; Lithospheric discontinuity; Continental evolution; Craton stabilization
1. Introduction Mechanisms tion of cratons research over 1988; Bickle,
controlling the origin and stabilizahave become the topic of vigorous the past 2 decades (Jordan, 1978, 1986; Vlaar, 1986; Helmstaedt and
Schulze, 1989; Abbott, 1991). Despite these efforts, progress in the understanding of continental evolution has been slow, owing in large part to difficulties in accessing the subcrustal lithosphere. Direct sampling is possible only through xenoliths erupted in kimberlite magmas and it is often unclear to what
0024-4937/99/$ - see front matter 9 1999 Elsevier Science B.V. All rights reserved. PII: S0024-4937(99)00020-1
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M.G. Bostock / Lithos 48 (1999) 1-16
extent these nodules are representative of mantle at a larger scale, and to what extent they have been modified in the eruption process (Harte and Hawkesworth, 1989). Geophysics (see also Jones, 1999), and seismology in particular, provides indirect means for assessing the larger-scale, present-day architecture of continental masses through remote sensing. Within seismology there exists an extensive arsenal of techniques for characterizing different styles of mantle structure. One category involves the identification of scattered waves which result from the interaction of a primary incident wave with a rapid change in elastic properties or density, here loosely termed discontinuity. Discontinuities in the upper mantle may arise through a variety of possible mechanisms. The Mohorovicic discontinuity (hereafter referred to as M) represents a compositional boundary and is a direct manifestation of planetary differentiation (Ringwood, 1975). Solid-solid phase changes in mantle minerals may result from increasing pressure with depth; phase changes in olivine are widely considered to be responsible for discontinuities at nominal depths of 410, 520, and 660 km depth within the mantle transition zone (Shearer, 1991). Rheology affords a third means of producing discontinuities in elastic properties, and one whose characterization yields important constraints on the internal tectonic fabric of continents (e.g., Leven et al., 1981; Karato, 1992; Gaherty and Jordan, 1995). The base of the continents is generally taken to represent a rheological boundary separating more rigid plate from more plastic convecting mantle, whereas rheological discontinuities within plates may reflect changes in deformation mechanism (Karato and Wu, 1993), or record the scars of ancient subduction, collision, and rifting. The object of this paper is to provide a short (and necessarily incomplete) review of seismological methods and research bearing on mantle lithospheric discontinuities, followed by a discussion of the collective implications of these studies for the understanding of continental evolution. It is written with the intention that it will provide non-seismologists engaged in the study of the continental mantle lithosphere with a better appreciation of both the constraints on lithospheric structure afforded by scattered seismic waves and their inherent limitations.
For a more general summary of seismic studies on upper mantle structure, the reader is referred to the work of Nolet et al. (1994).
2. Seismological considerations Seismological methods applied to the delineation of Earth structure can, in many instances, be classified in one of two broad categories: (i) transmission or time-of-flight approaches which consider the travel time required for a primary wave to propagate from source to receiver; and (ii) analyses of scattered waves which rely on the identification of secondary waves generated through the interaction of a primary wave (i.e., generated at a source) with discontinuities in physical properties within the Earth. Travel time tomography and analysis of surface wave dispersion are examples of the former category whereas seismic reflection profiling as used, for example, to delineate hydrocarbon traps in petroleum prospecting is perhaps the most obvious example of the latter. We will confine our attention primarily to this second category, though it should be emphasized that early contributions to our knowledge of continental lithospheric structure were derived in large part from studies of surface wave dispersion (e.g., Brune and Dorman, 1963; Dziewonski, 1971; Knopoff, 1972), and body wave delay times (e.g., Doyle and Hales, 1967; Sipkin and Jordan, 1976; Wickens and Buchbinder, 1980). Indeed, surface wave (e.g., Zielhuis and van der Hilst, 1996; EkstriSm et al., 1997; van der Lee and Nolet, 1997) and body wave (e.g., Grand, 1994; Humphreys and Dueker, 1994; VanDecar et al., 1995) tomography continue to play integral roles in the study of the continental mantle. In all seismological approaches, the detail with which Earth structure can be resolved is dependent upon the dominant wavelength of the interrogating seismic phase; however, techniques which exploit scattered phases hold an important advantage over transmission approaches in the retrieval of fine-scale structure. In transmission approaches the resolving power of the wave is governed by its Fresnel zone which represents the volume about the raypath between source and receiver which contributes signifi-
M.G. Bostock / Lithos 48 (1999) 1 - 16
cantly to the recorded seismogram. The diameter d of the Fresnel zone can be approximated by
d = v/AL/2 where L is the total path length traveled and A is dominant wavelength (related to frequency f as A = v / f where v is velocity). As an example, a teleseismic P-wave of frequency 1 Hz traveling at an average velocity of 10 k m / s along a total path length of 10 000 km possesses a Fresnel zone of d ~ 220 km. The implication is that structures existing within this volume are effectively averaged by the transmitted wave. Scattered phases, in contrast, are produced by interactions with structures at length scales comparable to the wavelength of the incident wave. For waves at near-vertical incidence in media which vary in only one coordinate, say depth, Richards (1972) has shown that significant reflections result from velocity gradients only if the width of the gradient zone is less than ~ A/4. This is a useful rule of thumb which has been applied, for example, by Benz and Vidale (1993) to constrain the transition thicknesses of the 410- and 660-km discontinuities to less than 4 km using ~ 1.0 Hz P'P' waves (contrast this value with the diameter of Fresnel zone in the example presented above). The Fresnel zone of reflected (and converted) waves must still be considered to assess the lateral extent of the discontinuity (or footprint) which is effectively sampled by the incident wave. It is therefore important in appreciating the results of a particular seismological study to consider the dominant wavelengths (or, equivalently, frequencies) which characterize the waves under observation. Transitions in physical properties which appear discontinuous to waves at longer wavelengths (or, equivalently, lower frequencies) may be relatively transparent to higher frequency phases. Conversely, higher frequency waves allow one to resolve complex structures in finer detail than long-period phases. The discussion of seismological methods that follows is thus largely based on a subdivision according to frequency.
3. Long-period seismology Our first topic for consideration concerns waveforms recorded on long-period seismometers. These
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instruments faithfully register signals in the frequency range between 0.05 and 0.010 Hz corresponding to wavelengths in the upper mantle of ~ 160-800 km for P-waves and ~ 90-600 km for S-waves. Long-period recordings are therefore useful for constraining average lithospheric properties at scales of major tectonic provinces through to continents. Shearer (1991) used digital long-period seismograms from a global data base to examine a variety of scattered phases including P-to-S (hereafter Ps) conversions, and precursors to the bottom-side reflections SS and PP (see Fig. l a,g for geometrical attributes of these phases) from the transition zone discontinuities. His analysis involved the stacking of large numbers of seismograms in bins of epicentral distance to enhance signal levels of the scattered phases. No clear evidence emerged for discontinuities at depths shallower than 410 km using either the global data set or a subset of seismograms with sampling restricted to continental platforms and shields. This result, as we discuss in a later section, poses certain restrictions on the nature of lithospheric discontinuities identified in more regional studies. Revenaugh and Jordan (1991) exploited another class of wave recorded on long-period seismograms to investigate upper mantle structure in a variety of tectonic environments in the western Pacific. ScS reverberations are transversely polarized waves which multiply (and perfectly) reflect between the Earth's free surface and the core mantle boundary, and hence decay in amplitude very slowly (see Fig. l b). With each circuit the ScS phase reflections interact with discontinuities in shear impedance (the product of shear velocity and density) to create a multiplicity of secondary scattered waves which can be processed to yield shear impedance profiles of the mantle. Their analysis of waves sampling continental regions revealed two lithospheric discontinuities termed H at ~ 6 0 km depth and L at ~ 2 2 0 km depth. The naming of these discontinuities derives from the individuals first responsible for proposing their existence, Lehmann (1955) and Hales (1969) and we will continue to use this nomenclature to refer to interfaces in other studies identified at or near similar levels. The presence of H on both continental and oceanic paths, an associated shear velocity increase of 3-4% and its location in depth led Revenaugh and
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M.G. Bostock/Lithos 48 (1999) 1-16
d) a) e)
b)
M
y V
g)
m
//// \\5
\
, \
\
\
\
Fig. 1. Schematic diagrams illustrating the geometrical raypaths represented by the variety of seismological studies discussed in the main text. Sources are marked as stars ( * ) and receivers as inverted triangles ( 9 P- and S-wave paths are shown as solid and dashed lines, respectively. Lithospheric discontinuity is outlined in bold. (a) Raypath corresponding to a PP precursor from underside reflection of lithospheric discontinuity. (b) Raypath for a single example of a ScS reverberation topside reflected from discontinuity. (c) Raypath for a Sp conversion. (d) Paths traveled by a series of rays interacting with a lithospheric discontinuity in a refraction profile. (e) Possible raypath of teleseismic Pn phase. (f) Near-vertical raypaths of reflections along a seismic reflection survey, g) Raypaths corresponding to near-receiver converted PD s (left) and SPD p phases (right).
Jordan (1991) to ascribe H to the phase transition from spinel to garnet lherzolite as proposed earlier by Hales (1969). In contrast to H, L generally appears only along corridors including some component of continental lithosphere and, further, exhibits an increase in depth towards older, more central portions of the Australian craton. This observation and the absence of a suitable candidate phase transition prompted Revenaugh and Jordan to postulate an anisotropic origin for L as the base of a layer below which aligned textures in peridotite are annealed. They remark on the coincidence of L with the depth at which deep seated kimberlite xenoliths are believed to originate. Gaherty and Jordan (1995) expanded on this idea in a comparison of diverse seismological data from oceanic and continental
(Australia) corridors in the southwest Pacific. They submitted evidence that L, as defined by long-period data, could be best explained as the base of an anisotropic lithosphere overlying more isotropic material embedded within the continental thermal boundary layer or tectosphere (Jordan, 1978, 1988). Revenaugh and Sipkin (1994) have applied the ScS migration technique to document discontinuity structure beneath China. Their interpretation of these data incorporated a more complex lithospheric structure involving the appearance of additional reflectors G, and U. G had previously been observed only on oceanic profiles (Revenaugh and Jordan, 1991) and represents a negative impedance contrast which varies regionally between 131 and 158 km depth. U occurs at similar depths to L (221-259 km) but was as-
M.G. Bostock/Lithos 48 (1999) 1-16
signed a different origin as the base of a low velocity zone associated with a downward decrease in strain heating. Long-period Sp phases are waves which result from mode conversion to P of an incident teleseismic S-wave at a discontinuity below a recording site (see Fig. l c). They are sensitive primarily to changes in shear velocity, and have been identified on longperiod seismograms in a number of studies concerning lithospheric structure. Jordan and Frazer (1975) noted evidence for H near 80 km depth below stations in eastern Canada. Baumgardt and Alexander (1984) exploited the availability of multiple stations in close proximity at the Large Aperture Seismic Array to identify the presence of a low velocity ( - 5 % ) channel between 171 and 226 km depth (L) below Montana. Despite clear Sp signals from the 410- and 660-km discontinuities, Bock (1988) was unable to find definitive evidence for a regionally coherent L-discontinuity beneath the Australian land mass. However, in an analysis of Sp phases recorded on the Baltic shield, Sacks et al. (1979) inferred the presence of a velocity inversion at approximately 250 km depth. On the basis of corroborative evidence from surface wave and heat flow studies, the authors attributed the discontinuity to the lithosphere-asthenosphere boundary.
4. Refraction seismology The results of deep sounding refraction profiles (DSPs) undertaken in the past 40 years have yielded important insights into lithospheric structure of the continental mantle. These have in general involved active sources (nuclear and chemical explosions), although in some circumstances earthquakes have also been employed (e.g., Hales et al., 1975; Leven et al., 1981; Walck, 1984). Such experiments include seismic energy to much higher frequencies (50-1 Hz, corresponding to P-wavelengths in the mantle lithosphere of 160 m - 8 km, respectively) than the studies outlined in Section 3 and thus permit the delineation of structure at considerably finer scales. The configuration of most DSPs has involved a linear array of densely spaced seismometers extending away from the source to at least 500 km distance to allow significant penetration below crustal levels
5
(Fig. l d). It is generally desirable to reverse the profile by positioning sources at both ends of the array to aid in distinguishing lateral variations from true 1D structure. In addition, more recent surveys have improved data coverage by including multiple shots along profile and employing a very dense sampling of seismometers (e.g., Henstock et al., 1998). The travel times of first and later arriving phases are measured and processed to produce a 1D (and, where data permit, 2D) velocity-depth profile characterizing the underlying crust and mantle. Although refraction seismology might more properly be classified a transmission approach according to our earlier definition, much important information regarding discontinuity structure arises from behavior of travel time branches corresponding to reflected waves; in particular, wide-angle reflections. Both reflected and refracted phases are most sensitive to structure near their turning points. Wide-angle reflections are larger amplitude than near-vertical incidence reflections, an attribute which facilitates identification in light of their arrival after the primary refracted energy. Refracted phases interact with horizontal structure at near-grazing angles of incidence, and under these circumstances, Richards' rule does not hold. Thus the sensitivity with which higher frequency refracted waves can discriminate between a first order discontinuity and a distributed gradient zone is somewhat reduced (e.g., Burdick and Orcutt, 1979). A general, though not universal, conclusion derived from long-range refraction profiles is that the continental mantle lithosphere (i.e., bounded above by M and below by the top of the asthenosphere or convecting mantle) may be divided into two depth regimes at a somewhat arbitrary level near 100-150 km. The results from a range of studies will be considered in light of this division. Project Early Rise (Iyer et al., 1969) represented one of the more important early DSPs and involved a series of linear arrays of seismometers radiating outwards from a shot point (using chemical explosives) in Lake Superior for distances of 2000-3000 km. Station spacing, typically of order 30 km, was sufficient to allow identification of the refracted branch from the 410-km discontinuity on several profiles. In addition, lithospheric mantle discontinuities were proposed at depths near 80-90 km (Green and Hales, 1968;
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M.G. Bostock/Lithos 48 (1999) 1-16
Mereu and Hunter, 1969) and near 125 km (Barr, 1967; Lewis and Meyer, 1968) along various profiles. These and other data prompted Hales (1969) and Hales et al. (1975) to suggest a common origin and an association of this discontinuity (H) with the phase change from spinel to garnet in aluminous peridotite. A wealth of information concerning upper mantle structure in central Asia has been assembled through the analysis of recordings from Peaceful Nuclear Explosions (PNE) made between 1960 and 1990 in the former Soviet Union. These explosions, comparable in size to medium-sized earthquakes (Mb 5.06.0), have been recorded on long (3000-4000 km) arrays of more closely spaced (to 10 km intervals) three-component seismometers, reversed, and augmented in profile with smaller chemical explosions. Early interpretations are summarized by Egorkin et al. (1987) and Pavlenkova (1996). Egorkin et al. (1987) analyzed a profile across northern Russia and noted the presence of strong lateral mantle heterogeneity at horizontal scales between 200 and 500 km, particularly within the depth range 80-100 km. Evidence was presented for both low-velocity waveguides and localized, anomalously high velocities below the Siberian craton above 100 km depth which did not appear to correlate with gravity anomalies. This latter observation has been cited as evidence for layers with well-developed elastic anisotropy in this region (Pavlenkova et al., 1996). A similar conclusion was reached by Fuchs (1979) in an early summary of constraints afforded on shallow lithospheric mantle structure from long-range refraction profiles. A number of more recent studies using PNE data (Ryberg et al., 1995, 1996; Tittgemeyer et al., 1996) have corroborated the results of earlier work and point to a number of other interesting features. The behavior of a phase known as teleseismic Pn is particularly intriguing. Although the precise mechanism of propagation is controversial (e.g., Menke and Richards, 1983), the Pn phase travels the majority of its path subhorizontally as a wave refracted (and possibly multiply reflected and/or scattered) in the shallow mantle below M (Fig. l e). Ryberg et al. (1995) noted that in central Asia a high frequency teleseismic Pn phase propagates efficiently over unusually large distances ( > 3000 kin) at group velocities of 8.1 k m / s and shows low spatial coherence at
10 km station intervals. They suggest that this feature develops through a laterally heterogeneous waveguide formed between M and a lower boundary near ~ 100 km depth. Tittgemeyer et al. (1996) demonstrate that it is possible to model many aspects of the Pn coda assuming a 1D model with random velocity fluctuations of +_4% superimposed on a positive velocity gradient in the top 100 km. The retrieval of greater detail in mantle structure may be achieved with still smaller station spacings. Hajnal et al. (1997) present three profiles, between 550 and 750 km in length, from the Trans-Hudson orogen in central North America exhibiting well-defined mantle discontinuities located at average depths of 75 and 158 km with some structural relief. These profiles are sampled at ~ 1 km intervals and, unlike the PNE sections, show laterally coherent, wide-angle reflections which the authors relate to alternating velocity layering below 100 km depth. In addition, these authors observe a high velocity (8.45 k m / s ) mantle block localized beneath a recently discovered kimberlite field, a scenario similar to that reported for the Siberian diamond provinces (Pavlenkova et al., 1996). Wide-angle reflections from shallow mantle interfaces near 50 km depth have been attributed to partial melt-induced low velocity zones beneath the Basin and Range province (Benz and McCarthy, 1994), and the western Canadian Cordillera (Clowes et al., 1995). The former study also identified a positive velocity jump near 80 km depth interpreted as H. Thybo and Perchuc (1997) examined a number of earthquake and active source refraction profiles from a range of continental environments and concluded that in most regions a low velocity zone beginning near a depth of 100 km underlies a generally stratified uppermost mantle. This interpretation is broadly consistent with the 1D and 2D interpretations made by Ryberg et al. (1995, 1996) of the PNE Quartz profile. Thybo and Perchuc (1997), following earlier speculation by Lehmann (1961) and Hales (1991), further suggest that the base of this low velocity zone coincides with the L discontinuity beneath continents and that it approaches the 410-km discontinuity in tectonically active regions. The existence of L was first ascertained on refraction profiles by Lehmann (1955, 1961) who analyzed travel times for P- and S-waves from earthquakes in North America
M.G. Bostock/Lithos 48 (1999) 1-16
and Europe. As noted in long-period studies, L appears to have continental affinity and has since been observed on refraction profiles from Australia (Hales et al., 1975; Leven et al., 1981; Drummond et al., 1982; Bowman and Kennett, 1990), and the former Soviet Union (Egorkin et al., 1987; Benz et al., 1992; Mechie et al., 1993; Priestly et al., 1994; Ryberg et al., 1996). In most cases L appears to be associated with a jump in P-wave velocity of between 2 and 4.5%. The observations in Australia were made using earthquake sources some of which occurred at depths to ~ 600 km. Deep earthquakes permit better constrained interpretation than possible using surface sources alone and allowed Leven et al. (1981) to associate L with a high velocity layer structure they interpreted to be near 30 km in thickness. Having eliminated phase transitions and compositional variations as likely explanations these authors preferred an interpretation in terms of anisotropy produced within a thin decoupling zone between plate and underlying asthenosphere. The preferred model in a reinterpretation (Priestly et al., 1994) of the PNE RIFT profile which traverses the Siberian platform also exhibits a layer-like structure with a velocity increase at 210 km followed by velocity reversal near 253 km depth. 1D models (Mechie et al., 1993) derived for the QUARTZ profile in central Siberia (subparallel to RIFT but displaced 1000 km to the west) suggest a more complicated layering exhibiting alternating velocities between 8.3 and 8.7 k m / s extending from 100 to 200 km depth. A subsequent 2D interpretation of the QUARTZ data by Ryberg et al. (1996) identifies a low velocity channel between 100 and 130 km below the northern portion of the profile, and a variable lithosphere/asthenosphere boundary ranging between 150 and 160 km in the south to 190-210 km in the north. Elements of this model are once more qualitatively similar to early work on PNE data from northern Siberia by Egorkin et al. (1987).
5. Reflection seismology Seismic reflection imaging has been extraordinarily successful in the detection of hydrocarbon traps. This success stems from a very dense sampling of the high frequency, near-vertical incidence, reflected P-wavefield (see Fig. lf) which allows for the stack-
7
ing of signals to reduce noise and the mapping of finer scale lateral variations in discontinuity structure than possible using other techniques. Mantle reflected P-waves are typically dominated by frequencies of the order of 20 Hz corresponding to wavelengths in the mantle of 400 m, and are sensitive to changes in "acoustic" impedance, the product of P-wave velocity and density. The improvement in resolution over other approaches, though costly, has led to the widespread application of seismic reflection methods in crustal profiling spearheaded by groups with such acronyms as COCORP, BIRPS, LITHOPROBE, and BABEL. Results from early work has indicated that the lithospheric mantle is largely transparent with comparatively low impedance contrasts, an observation underscored by the fact that M is often defined by the cessation of reflectivity with depth (see e.g., Best, 1991). Work in the past decade has revealed that, in some circumstances, it is possible to identify isolated mantle reflectors depths of 100 km or more. Perhaps the first, and certainly the most widely studied, discontinuities observed within the continental mantle using seismic reflection techniques are the Flanan thrust and so-called W reflectors below the North Sea, north of Scotland (Smythe et al., 1982). Their discovery has led to numerous subsequent surveys both in that region (McGeary and Warner, 1985; Warner and McGeary, 1987; Morgan et al., 1994; McBride et al., 1995; McBride et al., 1996; Warner et al., 1996) and nearby Scandinavia (BABEL Working Group, 1990; Lie and Husebye, 1994; Lie et al., 1990). Although alternate explanations involving rift-related extensional shearing within the mantle have been proposed (e.g., Reston, 1990; McBride et al., 1995), the most recent analyses based on detailed waveform modeling suggest that the Scottish reflectors represent high velocity ( > 8.4 kin/s), high density ( > 3.5 g / c m 3) layers of thickness greater than 2 km (probably 6-8 km thickness) within otherwise typical mantle (Warner et al., 1996). These characteristics clearly favor an interpretation in terms of subduction thrust faults which on the basis of surface deformation patterns are likely Precambrian in origin, with possible reactivation in more recent times (McBride et al., 1996). A number of other profiles have delineated similar features which can be tied directly to surface
8
M.G. Bostock / Lithos 48 (1999) 1 - 16
geology. Profiles shot in the Gulf of Bothnia (BABEL Working Group, 1990) reveal a thrust-like feature extending from the lower crust into the mantle which can be directly tied to the presence of an overlying 1.89 Ga volcanic arc complex. This provides evidence for plate convergence, subduction and accretion in the Early Proterozoic. In similar fashion,
Cook et al. (1997) have traced an interpreted Proterozoic subduction thrust from the lower crust into the mantle to depths of ~ 100 km below a 1.86-1.9 Ga volcanic arc in northwestern Canada (see Fig. 2). This study has also documented a horizontal reflector at ~ 80 km depth beneath the adjacent Archean Slave craton with a lateral extent likely exceeding
Fig. 2. Interpreted reflection seismic section from northwestern Canada (Cook et al., 1997) shown merged with a model of mantle stratigraphy below Yellowknife region (YKA) in the Slave province as determined from Ps converted phases (Bostock, 1998). Note presence of two sets of mantle reflections, one merging into lower crust at left, possibly continuous with feature L on teleseismic section; and a second near-horizontal set underlying much of the western Slave province which the teleseismic observations identify as 10 km thick layer H. See text for more discussion.
M.G. Bostock/Lithos 48 (1999) 1-16
200 km. Perhaps the most definitive evidence to date for subduction occurring in Archean time comes from the Abitibi granite-greenstone terrane of the Superior craton in eastern Canada. Calvert et al. (1995) interpret dipping reflections that extend 30 km into the mantle as a subduction thrust that led to a 2.69-Ga continental suture. Reflection signatures of very deep structures near 100-110 km depth between Denmark and Norway (Lie et al., 1990) and below the Urals (Knapp et al., 1996) at 130-170 km depth have been interpreted as marking the onset of the ductile asthenosphere.
6. Broadband seismology Broadband seismology refers to the study of seismograms recorded on a new generation of seismographs capable of recording signals between 50 and 0.01 Hz. Wide band width permits the analysis of a variety of seismic phases which interact with lithospheric discontinuities, some of which have been discussed in preceding sections. Ps conversions represent one class of phase with considerable potential for unveiling mantle structure. These are steeply propagating waves (see Fig. l g), generated by earthquakes in the distance range 30-100 ~ which are converted from P to S at a discontinuity in shear velocity below a three-component recording station. Typical frequencies of teleseismic P-waves between 5.0 Hz and 0.02 Hz imply wavelengths in the mantle between 1.6 km and 400 km; thus broadband recordings of Ps phases neatly bracket the frequency gap between seismic refraction and long-period studies. Although Ps conversions have been extensively used to investigate crustal (e.g., Langston, 1979) and transition zone (e.g., Vinnik, 1977; Gurrola et al., 1994) structure, their application to the lithospheric mantle has been largely confined to a few studies in the vicinity of active subduction zones (e.g., Crosson and Owens, 1987; Cassidy and Ellis, 1993). This owes in part to the interference of free surface multiples from M which arrive over a time window corresponding to that for direct conversions from the lithospheric mantle (Bostock, 1996), and are analogous to free surface multiples in reflection seismics. A recent study (Bostock, 1997, 1998) has shown that in areas where variations in crustal structure (in
9
particular topography on M) are small, a window into the lithospheric mantle devoid of multiples may be afforded by transverse component seismograms. In a laterally homogeneous, isotropic Earth no signal would be expected on the transverse component immediately following the first arriving P-wave. However, at a site on the southern part of the Slave craton, prominent and laterally coherent arrivals are evident and signal the presence of a well-developed anisotropic mantle stratigraphy (see Fig. 2). The layering occurs in three sequences at nominal depths of 75 km (H), 135 km (termed X), and ~ 190 km (L). The shallowest of these is best defined and comprises a near horizontal ~ 10 km thick layer exhibiting _+5% anisotropy. Consequently, the layer appears to be high velocity from some azimuths and low velocity from others relative to ambient mantle. The upper interface of the layer is sharp to the highest frequencies in the data (5.0 Hz) suggesting a transition of less than 400 m whereas the lower boundary is more diffuse and likely extends over an interval > 2 km in width. The deeper discontinuities are more difficult to characterize; however, X appears to represent a sequence of at least two layers of order 5-10 km thickness while L is a simpler, single interface which dips from 170 km in the west to 230 km in the east over a lateral distance of ~ 120 km (Fig. 2). On the basis of petrographic evidence from kimberlite xenoliths and results from the coincident reflection profile (Cook et al., 1997), Bostock (1998) interpreted these features as related to multiple episodes of subduction underplating extending back to Archean time. A preliminary analysis of data from other broadband stations on the Canadian land mass provides evidence for similar lithospheric mantle layering at other sites, in particular within Precambrian settings. These features often occur as single layers ~ 10 km in thickness within the top 100 km of mantle (Fig. 3). A number of scattered phases associated with incident S-waves and recorded at intermediate periods are potentially useful in the delineation of lithospheric discontinuities. Zandt and Randall (1985) observed post-critically reflected, shear-coupled Pwaves (known as SPD p) from a shallow (70-80 km) high velocity contrast discontinuity at several stations in North America which they interpreted as H. Similar observations of SPD p phases have been made
10
M.G. Bostock / Lithos 48 (1999) 1 - 16 i
i
i
i
1
i
i
structures. Such arrays will permit the application of processing procedures similar to those used in reflection seismics to migrate scattered waves into images of mantle heterogeneity. A demonstration of the potential of such an approach is provided by Revenaugh (1995) who used some 5000 short-period recordings of teleseismic events from the Southern California Seismic Network to map the southern flank of subducting lithosphere beneath the Transverse Ranges.
7. Discussion
1 0
50 160 l go 200 250 3bo ago 4oo Depth [kin]
Fig. 3. Mantle reflectivity as a function of depth determined from teleseismic PDs phases at a selection of stations situated on the Canadian Shield. These profiles are sensitive to discontinuities in anisotropic material and thus do not generally show a strong M signature. Profiles are derived from transverse component seismograms over a narrow range of back-azimuths (BAZ). Arrows point to prominent discontinuities within the mantle lithosphere. The bottom profile (YKA) corresponds to the teleseismic section shown in Fig. 2.
by Clarke and Silver (1991) and Earle et al. (1994). In the vicinity of subduction zones, P precursors to near-vertical incidence ScS, known as ScSp, have proven useful in characterizing the slab-mantle interface (see e.g., Snoke et al., 1977; Nakanishi et al., 1981; Helffrich and Stein, 1993); however, more general application to the interior of continents is hampered by a general paucity of earthquakes away from plate boundaries. Sp phases discussed in an earlier section do not suffer in this regard and may be used in a similar fashion (Farra et al., 1991). Finally, it is worth remarking upon the future use of large dense arrays of broadband seismometers for the delineation of detailed and complex lithospheric
There are a number of possibilities to consider in placing the varied observations outlined above into a global context. A reasonable approach is to attempt to relate them directly to the essential precepts of plate tectonics and the Wilson cycle, specifically the processes of subduction, collision and rifting. Several reflection profiles (Smythe et al., 1982; Calvert et al., 1995; Cook et al., 1997) provide clear evidence of mantle discontinuities merging into lower continental crust of Precambrian age. Similarities between these reflectors and the seismic signatures characterizing present-day convergent margins (e.g., Fuis and Clowes, 1993) together with the context provided through surface geology render it difficult to entertain mechanisms which are not subduction related. The most obvious candidates are major subduction thrust faults whose activity has terminated in continental collision. It appears likely, therefore, that subduction thrust faults are capable of creating seismic discontinuities, both compositional and rheological in nature, preserved over geologic time. Moreover, it is reasonable to conjecture that at least some proportion of mantle discontinuities observed in other studies, where direct association with the crust is not in evidence, bear some immediate relation to emplacement of subducted oceanic lithosphere. Two studies discussed in previous sections provide less direct but nonetheless strong supporting evidence for ancient oceanic crust stranded in the subcontinental lithosphere where an explicit retracing to a suture or overlying arc cannot be made. Warner et al. (1996) and Bostock (1998) both note the presence of subhorizontal layers ~ 10 km in
M.G. Bostock/Lithos 48 (1999) 1-16
thickness within the shallow mantle, possessing sharp upper and more diffuse lower boundaries. In the former study, reflection coefficients and gravity profiles of the North Sea W reflector suggest that the layer most likely comprises eclogitized oceanic crust. A similar conclusion is reached by Bostock (1998) for the H layer (Fig. 2) below the Archean Slave province on the basis of a strongly anisotropic signature coupled with information from xenolith studies and seismic reflection profiling. In both cases structural relationships suggest a Precambrian origin, and in the latter instance the age may well be early Archean. These interpretations combined with the observed lateral extent of the layers ( > 100-200 km) and the absence of significant dip implies that their original emplacement involved a shallow mode of subduction. This latter observation bears important implications for our understanding of continental evolution by providing a mechanism to account for several fundamental geological and geodynamic observations which distinguish Archean cratons from continental material formed in more recent times (see e.g., Abbott, 1991). Although shallow subduction is not widely represented among the range of presentday convergent margins (Sacks, 1983; Abbott et al., 1994b), there are several lines of evidence that point to greater incidence in the past. Arguments have been made that increased heat flow in Archean time was accommodated through greater ridge length with a less significant role played by increased oceanic crustal thickness (Abbott and Menke, 1990; Sleep, 1992). The net buoyancy of oceanic lithosphere under these conditions should be increased over that of the present day through (i) a younger mean age of plate at subduction, and (ii) a decrease in content of iron and garnet within the mantle lithospheric column (Jordan, 1979) accruing from higher melting temperatures (Sleep and Windley, 1982; Bickle, 1986; Abbott et al., 1994a). These arguments have led several authors to suggest that shallow subduction figured more prominently in Archean tectonics than it does at present (Vlaar, 1986; Abbott, 1991; Abbott et al., 1994b). A number of studies have emphasized the importance of aseismic ridges and oceanic plateaus in initiating shallow subduction in the Phanerozoic (Livaccari et al., 1981; Pilger, 1981; Henderson et al., 1984) leading to the proposal that
11
similar but more widespread processes facilitated the stabilization of large tracts of continental lithosphere in the Archean (de Wit et al., 1992; Abbott and Mooney, 1995). The hypothesis that shallow subduction is responsible, at least in part, for craton stabilization can be tested through its consistency with other constraints afforded by seismic studies. First, consider that shield regions of North America, Eurasia, and western Australia are characterized by surface geology that reflects an amalgamation of integral blocks (portions of former plates) with typical scale lengths of 300500 km brought together in the terminal phases of convergence and continental collision (e.g., Hoffman, 1989). A dominantly horizontal stratification with similar lateral wavelengths should be expected to characterize the underlying lithospheric mantle if these blocks had originally stabilized through processes involving shallow subduction. The details of this stratification should clearly depend on the individual tectonic history experienced by a particular block. Older, Archean blocks might record several episodes of rapid, shallow subduction allowing for the development of pronounced low temperature keels capable of harboring diamonds soon after stabilization of overlying crust (Richardson et al., 1984). Stabilization of younger blocks would be accommodated by fewer (or single) subduction episodes as the incidence of shallow subduction waned through the Proterozoic (Abbott et al., 1994a). This expected variation in character of mantle stratigraphy from province to province is generally consistent with the variability of depths at which lithospheric interfaces have been identified in seismic refraction and reflection studies (see preceding sections), and the abrupt lateral transitions in mantle velocity structure documented, for example, in western-central North America (Henstock et al., 1998). It also explains the lack of coherent energy from levels shallower than 400 km in the global waveform stacks of Shearer (1991). The degree of horizontal stratification, averaged laterally, would moreover tend to be strongest at the shallower levels of the mantle, and diminish at greater depths. Such a model is consistent with stochastic models of mantle velocity heterogeneity designed by Kennett and Bowman (1990) (see also Jordan et al., 1996) to satisfy both short and longperiod seismic observations from the continents. Fur-
12
M.G. Bostock / Lithos 48 (1999) 1-16
ther, note that by laterally juxtaposing blocks exhibiting broadly similar but individually distinct horizontal stratification concentrated at shallower levels, it is possible to conceive of an imperfect, scattering waveguide capable of reproducing the essential characteristics of the teleseismic Pn phase documented by Ryberg et al. (1995). Fig. 4 shows a schematic realization of such a structure based partly on the lithospheric cross-section for northwestern Canada shown in Fig. 2. At some depth this shallow, dominantly horizontal stratification must yield to mantle with a less organized fabric. The depth at which this occurs and the transition to convecting mantle will depend on the complex interplay of a number of factors which include the age and history of constituent blocks, topology of shallow subduction (Sacks, 1983), rheology of the lithospheric column as dictated by thermal structure (Karato and Wu, 1993) and composition, and the nature of slab break-off (Davies and von Blanckenburg, 1995). It will likely be manifest by a change in the character of seismic anisotropy, consistent with the results of Gaherty and Jordan (1995) who identify the layer L from long-period ScS studies with an extinction of anisotropy below ~ 200 km depth in Australia. A similar result is obtained by Montagner and Tanimoto (1991) who examined long-period surface waves in a more global study.
~
'..,'
Effective seismic anisotropy, especially at longer periods, may result from macroscopic foliation/lineation at a range of scales and involving a variety of lithologies (Fuchs, 1979; Bostock, 1998), in addition to the oft-cited lattice preferred orientation of olivine (Karato and Wu, 1993). The presence of a significant component of azimuthal anisotropy may, as an alternative or supplement to lateral variability, aid in explaining the absence of coherent discontinuity structure above 410 km in the global, long-period waveform stacks of Shearer (1991). The foregoing model of continental stabilization via shallow subduction is acknowledged to be simplistic and imprecisely specified, and though capable of explaining a range of seismic observations concerning lithospheric discontinuities, it also raises a host of new questions. One example concerns the nature of H observed at long period (Revenaugh and Jordan, 1991); if shallow discontinuities in the mantle are properly ascribed to stranded oceanic lithosphere, then what is the nature of H observed below ocean basins? A second example concerns the dynamics of shallow subduction. It is widely accepted that the Laramide orogeny in the western United States involved large-scale shallow subduction (Livaccari et al., 1981; Bird, 1988) and culminated in equally extensive lithospheric delamination (Bird, 1988; Humphreys, 1995). The final stages of this
,'"
Fig. 4. Schematic diagram, loosely based on information in Fig. 2, depicting continental lithosphere as a juxtaposition of blocks stabilized through shallow subduction and brought together through continental collision.
M.G. Bostock/Lithos 48 (1999) 1-16
latter process are captured in the images of upper mantle shear velocity derived by van der Lee and Nolet (1997) from inversion of surface wave waveforms. Although other mechanisms for lithospheric delamination have been proposed (Houseman et al., 1981), it would appear plausible that the small buoyancy excess responsible for shallow subduction (acquired, say, in the form of an oceanic plateau or aseismic ridge) is subsequently cancelled through the transformation of subducted basaltic crust into eclogite (Cloos, 1993) with an accompanying density increase of 15% (Ahrens and Schubert, 1975). How shallow subduction has thus led to delamination in the Phanerozoic setting and stabilization in the Archean (Abbott, 1991; Hart et al., 1997; Bostock, 1998) is another outstanding question that will require input from geodynamics and petrology. Seismology will undoubtedly continue to play an important role in addressing these issues and the more general conundrum of continental evolution through increasingly detailed mapping of lithospheric discontinuity structure.
Acknowledgements I am indebted to Fred Cook for permission to use the interpreted LITHOPROBE reflection profile in Fig. 2. Constructive comments made on an early version of the manuscript by Ron Clowes and Andy Frederiksen are gratefully acknowledged. Rob van der Hilst and an anonymous referee are thanked for helpful reviews. This research was supported through a Natural Sciences and Engineering Research Council of Canada collaborative grant (NSERC CPG0163720), and I express my gratitude to group members Don Francis, Herb Helmstaedt, Shao-cheng Ji, Alan Jones, John Ludden, and Marianne Mareschal (deceased) for fruitful and engaging discussions.
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Rev. Geophys. Space Phys., Natl. Rep. IUGG Suppl., pp. 231-242. Abbott, D., Burgess, L., Longhi, J., Smith, W.H.F., 1994a. An empirical thermal history of the Earth's upper mantle. J. Geophys. Res. 99, 13835-13850. Abbott, D., Drury, R., Smith, W.H.F., 1994b. Flat to steep transition in subduction style. Geology 22, 937-940. Ahrens, T.J., Schubert, G., 1975. Gabbro-eclogite reaction rate and its geophysical significance. Rev. Geophys. Space Phys. 13, 383-400. BABEL Working Group, 1990. Evidence for early Proterozoic plate tectonics from seismic reflection profiles in the Baltic shield. Nature 348, 34-38. Barr, K.G., 1967. Upper mantle structure in Canada from seismic observations using chemical explosions. Can. J. Earth Sci. 4, 961-975. Baumgardt, D.R., Alexander, S.S., 1984. Structure of the mantle beneath the Montana LASA from analysis of long-period, mode-converted phases. Bull. Seismol. Soc. Am. 74, 16831702. Benz, H.M., McCarthy, J., 1994. Evidence for an upper mantle low velocity zone beneath the southern Basin and Range-Colorado Plateau transition zone. Geophys. Res. Lett. 21, 509512. Benz, H.M., Vidale, J.E., 1993. Sharpness of upper-mantle discontinuities determined from high-frequency reflections. Nature 365, 147-150. Benz, H.M., Unger, J.D., Leith, W.S., Mooney, W.D., Solodilov, L., Egorkin, A.V., Ryaboy, V.Z., 1992. Deep seismic sounding in northern Eurasia. EOS 73, 297-300. Best, J.A., 1991. Mantle reflections beneath the Montana Great Plains on Consortium for Continental Reflection Profiling Seismic Reflection data. J. Geophys. Res. 96, 4279-4288. Bickle, M.J., 1986. Implications of melting for stabilization of the lithosphere and heat loss in the Archaean, Earth Planet. Sci. Lett. 80, 314-324. Bird, P., 1988. Formation of the Rocky Mountains, western United States: a continuuum computer model. Science 239, 1501-1507. Bock, G., 1988. Sp phases from the Australian upper mantle. Geophys. J. 94, 73-81. Bostock, M.G., 1996. A seismic image of the upper mantle beneath the North American craton. Geophys. Res. Lett. 23, 1593-1596. Bostock, M.G., 1997. Anisotropic upper mantle stratigraphy and architecture of the Slave craton. Nature 390, 392-395. Bostock, M.G., 1998. Mantle stratigraphy and evolution of the Slave province. J. Geophys. Res. 103, 21183-21200. Bowman, J.R., Kennett, B.L.N., 1990. An investigation of the upper mantle beneath NW Australia using a hybrid seismograph array. Geophys. J. Int. 101,411-424. Brune, J., Dorman, J., 1963. Seismic waves and Earth structure in the Canadian Shield. Bull. Seismol. Soc. Am. 53, 167-210. Burdick, L.J., Orcutt, J.A., 1979. A comparison of the generalized ray and reflectivity methods of waveform synthesis. Geophys. J. R. Astron. Soc. 58, 261-278. Calvert, A.J., Sawyer, E.W., Davis, W.J., Ludden, J.N., 1995.
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Archaean subduction inferred from seismic images of a mantle suture in the Superior Province. Nature 375, 670-673. Cassidy, J.F., Ellis, R.M., 1993. S-wave velocity structure of the northern Cascadia subduction zone. J. Geophys. Res. 98, 4407-4421. Clarke, T.J., Silver, P.G., 1991. A procedure for the systematic interpretation of body wave seismograms: I. Application to Moho depth and crustal properties. Geophys. J. Int. 104, 41-72. Cloos, M., 1993. Lithospheric buoyancy and collisional orogenesis: subduction of oceanic plateaus, continental margins, island arcs, spreading ridges, and seamounts. Geol. Soc. Am. Bull. 105, 715-737. Clowes, R.M., Zelt, C.A., Amor, J.R., Ellis, R.M., 1995. Lithospheric structure in the southern Canadian Cordillera from a network of seismic refraction lines. Can. J. Earth Sci. 32, 1485-1513. Cook, F.A., Van der Velden, A.J., Hall, K.W., 1997. Upper mantle reflectors beneath the SNORCLE transect images of the base of the lithosphere? In: Cook, F., Erdmer, P. (Eds.), S l a v e - N o r t h e r n Cordillera Lithospheric Evolution (SNORCLE) Transect and Cordilleran Tectonics Workshop Meeting. Lithoprobe Report No. 56. University of Calgary, pp. 58-62. Crosson, R.S., Owens, T.J., 1987. Slab geometry of the Cascadia subduction zone beneath Washington from earthquake hypocenters and teleseismic converted waves. Geophys. Res. Lett. 14, 824-827. Davies, J.H., von Blanckenburg, F., 1995. Slab breakoff: a model of lithosphere detachment and its test in the magmatism and deformation of collisional orogens. Earth Planet. Sci. Lett. 129, 85-102. de Wit, M.J., Roering, C., Hart, R.J., Armstrong, R.A., de Ronde, C.E.J., Green, R.W.E., Tredoux, M., Peberdy, E., Hart, R.A., 1992. Formation of an Archaean continent. Nature 357, 553562. Doyle, H.A., Hales, A.L., 1967. An analysis of the travel times of S-waves to North American stations in the distance range 280-82 ~ Bull. Seismol. Soc. Am. 57, 761-771. Drummond, B.J., Muirhead, K.J., Hales, A.L., 1982. Evidence for a seismic discontinuity near 200 km depth under a continental margin. Geophys. J. R. Astron. Soc. 70, 67-77. Dziewonski, A.M., 1971. Upper mantle models from 'pure path' dispersion data. J. Geophys. Res. 76, 2587-2601. Earle, P.S., Shearer, P., Zandt, G., Walter, W., 1994. A global survey for the 80-km discontinuity utilizing three-component broadband data. EOS 75, 435. Egorkin, A.V., Zuganov, S.K., Pavlenkova, N.A., Chernyshev, N.M., 1987. Results of lithospheric studies from long-range profiles in Siberia. Tectonophysics 140, 29-47. EkstriSm, G., Tromp, J., Larson, E.W.F., 1997. Measurements and global models of surface wave propagation. J. Geophys. Res. 102, 8137-8157. Farra, V., Vinnik, L.P., Romanowicz, B., Kosarev, G.L., Kind, R., 1991. Inversion of teleseismic S particle motion for azimuthal anisotropy in the upper mantle: a feasibility study. Geophys. J. Int. 106, 421-431.
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reverberations: 3. The upper mantle. J. Geophys. Res. 96, 19781-19810. Revenaugh, J., Sipkin, S.A., 1994. Mantle discontinuity structure beneath China. J. Geophys. Res. 99, 21911-21927. Richards, P.G., 1972. Seismic waves reflected from velocity gradient anomalies within the Earth's upper mantle. J. Geophys. 38, 517-527. Richardson, S.H., Gurney, J.J., Erlank, A.J., Harris, J.W., 1984. Origin of diamonds in old enriched mantle. Nature 310, 198-202. Ringwood, A.E., 1975. Composition and Petrology of the Earth's Mantle. McGraw-Hill, New York. Ryberg, T., Fuchs, K., Egorkin, A.V., Solodilov, L., 1995. Observation of high-frequency teleseismic Pn on the long-range Quartz profile across northern Eurasia. J. Geophys. Res. 100, 18151-18163. Ryberg, T., Wenzel, F., Mechie, J., Egorkin, A.V., Fuchs, K., Solodilov, L., 1996. Two-dimensional velocity structure beneath Northern Eurasia derived from the super long-range seismic profile Quartz. Bull. Seismol. Soc. Am. 86, 857-867. Sacks, I.S., 1983. The subduction of young lithosphere. J. Geophys. Res. 88, 3355-3366. Sacks, I.S., Snoke, J.A., Husebye, E.S., 1979. Lithosphere thickness beneath the Baltic shield. Tectonophysics 56, 101-110. Shearer, P.M., 1991. Constraints on upper mantle discontinuities from observations of long-period reflected and converted phases. J. Geophys. Res. 96, 18147-18182. Sipkin, S.A., Jordan, T.H., 1976. Lateral heterogeneity of the upper mantle determined from the traveltimes of multiple-ScS. J. Geophys. Res. 81, 6307-6320. Sleep, N.M., 1992. Archean plate tectonics: what can be learned from continental geology?. Can. J. Earth Sci. 29, 2066-2071. Sleep, N.M., Windley, B.F., 1982. Archean plate tectonics: constraints and inferences. J. Geol. 90, 363-379. Smythe, D.K., Dobinson, A., McQuillin, R., Brewer, J.A., Matthews, D.H., Blundell, D.J., Kelk, B., 1982. Deep structure of the Scottish Caledonides by the MOIST reflection profile. Nature 299, 338-340.
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LITHOS ELSEVIER
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Lithos 48 (1999) 17-43
The deep structure of the Australian continent from surface wave tomography Frederik J. Simons *, Alet Zielhuis, Rob D. van der Hilst Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Rm 54-517A, Cambridge, MA 02139, USA Received 16 April 1999; received in revised form 10 May 1999; accepted 11 May 1999
Abstract We present a new model of 3-D variations of shear wave speed in the Australian upper mantle, obtained from the dispersion of fundamental and higher-mode surface waves. We used nearly 1600 Rayleigh wave data from the portable arrays of the SKIPPY project and from permanent stations (from AGSO, IRIS and GEOSCOPE). AGSO data have not been used before and provide better data coverage of the Archean cratons in western Australia. Compared to previous studies we improved the vertical parameterization, the weighting scheme that accounts for variations in data quality and reduced the influence of epicenter mislocation on velocity structure. The dense sampling by seismic waves provides for unprecedented resolution of continental structure, but the wave speed beneath westernmost Australia is not well constrained. Global compilations of geological and seismological data (using regionalizations based on tectonic behavior or crustal age) suggest a correlation between crustal age and the thickness and composition of the continental lithosphere. However, the age and the tectonic history of crustal elements vary on wavelengths much smaller than have been resolved with global seismological studies. Using our regional upper mantle model we investigate how the seismic signature of tectonic units changes with increasing depth. At large wavelengths, and to a depth of about 200 km, the inferred velocity anomalies corroborate the global pattern and display a progression of wave speed with crustal age: slow wave propagation prevails beneath the Paleozoic fold belts in eastern Australia and wave speeds increase westward across the Proterozoic and reach a maximum in the Archean cratons. The high wave speeds associated with Precambrian shields extend beyond the Tasman Line, which marks the eastern limit of Proterozoic outcrop. This suggests that parts of the Paleozoic fold belts are underlain by Proterozoic lithosphere. We also infer that the North Australia craton extends off-shore into Papua New Guinea and beneath the Indian Ocean. For depths in excess of 200 km a regionalization with smaller units reveals that some tectonic subregions of Proterozoic age are marked by pronounced velocity highs to depths exceeding 300 km, but others do not and, surprisingly, the Archean units do not seem to be marked by such a thick high wave speed structure either. The Precambrian cratons that lack a thick high wave speed " k e e l " are located near passive margins, suggesting that convective processes associated with continental break-up may have destroyed a once present tectosphere. Our study suggests that deep lithospheric structure varies as much within domains of similar crustal age as between units of different ages, which hampers attempts to find a unifying relationship between seismic signature and lithospheric age. 9 1999 Elsevier Science B.V. All fights reserved. Keywords: Continental lithosphere; Rayleigh wave; SKIPPY project; Broadband data; Waveform tomography; Australia
* Corresponding author. Tel.: + 1-617-253-0741; fax: + 1-617-258-9697; E-mail:
[email protected] 0024-4937/99/$ - see front matter 9 1999 Elsevier Science B.V. All fights reserved. PII: S 0 0 2 4 - 4 9 3 7 ( 9 9 ) 0 0 0 4 1 - 9
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F.J. Simons et al./Lithos 48 (1999) 17-.43
1. Introduction
The inference that continental geotherms in the Archean were not much different from the averages seen today, despite pervasively (200-300 ~) hotter mantle temperatures, has been supported (but also refuted, e.g., Strong and Stevens, 1974) by both observations (Burke and Kind, 1978; Boyd and Gurney, 1986; Boyd, 1989) and numerical simulations (Richter, 1984, 1985; Lenardic, 1998) (see also Jaupart and Mareschal, 1999, and Nyblade, 1999, this volume). Combined with experimental data on diamond stability in the mantle, the age of the continental lithospheric mantle (CLM) inferred from diamond inclusions (Richardson et al., 1984) and Re-OS dating of mantle xenoliths (Pearson et al., 1995; Pearson, 1999, this volume), and the correlation between composition of continental mantle and the age of the overlying crust (Griffin et al., 1998) this suggests the presence of a thick (> 175 km) Archean and EarlyProterozoic subcontinental lithosphere that has stabilized shortly after being formed and has remained coupled to the crust ever since. This "cold" lithosphere, with geothermal gradients generally less than 23 ~ km -1 (Burke and Kind, 1978), must have been stabilized against convective disruption by its composition and its rheological properties (Jordan, 1988; Richter, 1988; De Smet et al., 1999, this volume; Shapiro et al., 1999a;, this volume). Cooling of the lithosphere provides a satisfactory explanation for the geophysical signature (e.g., gravity, bathymetry, seismic properties) of oceanic plates (Sclater et al., 1981), but the large CLM thickness inferred from seismic imaging cannot be explained by cooling alone. If a cold, negatively buoyant continental lithosphere is stabilized by rheology (that is, the strength of the plate) the thermally induced density contrasts would produce anomalies in the longwavelength gravity field and geoid, which are not observed (Kaula, 1967; Jordan, 1975b, 1978; Shapiro et al., 1999b, this volume). Considering geological, geochemical, and geophysical evidence, Jordan (1975b, 1978) postulated a mechanism for producing a petrologically distinct chemical boundary layer under the cratonic parts of the continents. In his isopycnic model the negative buoyancy of cold lithosphere is compensated by positive compositional buoyancy. Neutrally buoyant and probably more viscous than
the surrounding mantle, this "keel" of depleted, refractory continental mantle, or tectosphere, is thought to remain coupled to the crust for billions of years and thus plays a crucial role in regulating continental development and stability (Jordan, 1981a, 1988). The presence of such "keels" has been corroborated by seismological studies that reveal a (300-500 km) thick, anomalously fast lithosphere under some of the oldest parts of the continents (Jordan, 1975a; Sipkin and Jordan, 1975, 1976; Woodhouse and Dziewonski, 1984; Suet al., 1994; Masters et al., 1996; EkstriSm et al., 1997; van der Lee and Nolet, 1997b). To first order this relationship seems to hold globally (Okal, 1977; Jordan, 1981b; Polet and Anderson, 1995), but, as we show in this study, significant departures may occur on smaller length scales. The lateral variation in upper mantle structure, such as the differences between oceanic and continental regions and between stable or tectonically active domains, can be studied with a variety of seismic imaging techniques. The type of data used depend on the specific research objectives. Elsewhere in this volume, Bostock (1999) reviews techniques used for high-resolution imaging of the crust and lithospheric mantle with phase conversions and scattering of seismic body waves. Here we consider variations in elastic properties on a larger scale than considered by Bostock; these can be constrained by a variety of methods and data, for example dispersion curves of fundamental-mode surface waves (Toks~Sz and Anderson, 1966; Priestley, 1999, this volume), the dispersion of singlets in the eigenfrequency multiplets of fundamental spheroidal and toroidal modes (Dahlen, 1976), vertical delay times of (multiply) reflected body wave phases (Sipkin and Jordan, 1975, 1976), and higher mode phase and group velocities (Cara, 1979). With the advent of different methods of surface-wave tomography (Woodhouse and Dziewonski, 1984; Nolet, 1990; Trampert and Woodhouse, 1995, 1996; EkstriSm et al., 1997) direct images have been made of the aspherical structure of the upper mantle. Typically, the horizontally propagating surface waves provide better radial resolution than vertically incident body waves, especially when higher modes are included in the analysis. Body waves potentially provide better lateral resolution of lithospheric structure. In general, global data sets
F.J. Simons et al./Lithos 48 (1999) 17-43
resolve structure on the scale of sub-continental age provinces. For example, the two-fold division of Australia into a Phanerozoic and Precambrian part is resolved in the global models (EkstriSm et al., 1997), but these do not constrain variations in structure over distances less than about 1000 km (Laske and Masters, 1996). Regional studies can provide better horizontal resolution, in particular if additional data from regional seismometer networks are available. Australia is well suited to investigate deep continental structure. The makeup of the continent is extremely varied (Fig. 1). Based on outcrop, continental Australia can be divided in a western, central, and eastern domain (see Fig. l c) but smaller-scale units have been identified as well (see Fig. l a, b). The western third of the continent comprises granites and greenstones of the Pilbara and Yilgarn blocks, which formed in the Archean (3500 and 3100 Ma, respectively) and have been stable since at least 2300 Ma (Plumb, 1979). There is little or no Phanerozoic cover of these cratonic units. Central Australia consists of a series of intracratonic Late-ProterozoicEarly-Paleozoic basins separated by fault-bounded blocks exposing Mid-Proterozoic basement rocks (Lambeck, 1983). The Tasman Line first drawn by Hill in 1951 (Veevers, 1984) but different interpretations exist separates this Precambrian outcrop from exposed Phanerozoic formations in the east. Its definition is based largely on surface geology and, in regions of limited exposure (such as across the Eromanga basin), on lineations in gravity and magnetic anomaly maps (Murray et al., 1989; Wellman, 1998). Moreover, the continent is favorably located with respect to zones of active seismicity (Fig. 2), which provide ample sources for seismic tomographic imaging, and the SKIPPY seismometry project (van der Hilst et al., 1994) has provided data coverage that allows surface wave tomography with unprecedented resolution. In this paper we briefly describe the tomographic technique that we have used to determine lateral variations in shear wave speed in the Australian mantle and discuss the relationship between the thickness of the high wave speed lithosphere and variations in crustal age. We present our model both in terms of the inversion results and as wave speed averages over well-defined geotectonic regions. Using regionalizations at different spatial scales we
19
show that the deep structure of the Australian continent varies significantly, not only across the largescale tectonic units but also within domains of roughly the same age. These observations are important for our understanding of the Australian continent but also have ramifications for studies of deep continental structure and evolution on the basis of global regionalizations of crustal age and tectonic history.
2. Imaging with seismic waves" seismic tomography Tomography is a technique for reconstructing a function ("the unknowns", or "the model") from projections ("the data") along a set of curves. This relationship is often expressed as an integral over a certain volume V, fv gi(r)6x(r) d r = bi, with gi(r) the Fr6chet derivative (or sensitivity kernel) describing the functional dependence of the measurements b i on the model perturbations 6 x(r) (or in linearized form as a system of normal equations Ax = b, with A the sensitivity matrix containing the appropriate Fr6chet derivatives and x and b the model and data vector, respectively). If the medium under study is the human body, then the function might be the density of organ tissue, and the data used to constrain it might be the intensity of transmitted X-rays (the principle behind many kinds of medical tomography (Herman, 1979)). For imaging the Earth, one uses seismic waves, which are affected by anomalous structure, so that the phase arrival time, amplitude, or entire waveform differ from the ones expected in a spherical reference Earth model. Such differences are then interpreted in terms of velocity and attenuation variations of seismic waves within the Earth. A major complication is that the sources (earthquakes or man-made explosion) and receivers (seismometers) are distributed very unevenly over the surface of the Earth so that some regions are constrained by many data whereas others are not sampled at all. This renders the tomographic problem underdetermined (that is, not all unknowns can be determined independently) and the solution is nonunique. Out of a large number of solutions we choose a solution by minimizing a penalty function,
20
F.J. Simons et a l . / Lithos 48 (1999) 17-43
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F.J. Simons et a l . / Lithos 48 (1999) 17-43 /
"x,
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Fig. 2. Locations of the SKIPPY, Iris, GEOSCOPE,and AGSO stations used in this study. The solid stars depict epicenters of all earthquakes (source: Engdahl et al., 1998) for which data were used in this study ( N = 336). Abbreviations used: QLD: Queensland, NSW: New South Wales, VIC = Victoria, SA = South Australia, TAS = Tasmania, NT = Northern Territory, WA = West Australia. Gray dashed lines depict plate boundaries.
which typically includes regularization terms (also known as damping) and accounts for a priori information. Different kinds of seismic tomography exist (see, for example, the overview by Nolet (1987)). The data that can be used differ in frequency content and in the way they sample Earth's interior (see Fig. 3). Body-wave tomography often makes use of travel-
time perturbations from reference times. With surface waves, which typically have lower frequencies than body waves, the principle is the same, but they sample the Earth in a different manner. Rather than having their sensitivity to Earth structure concentrated along a ray path (such as in Fig. 3b) this sensitivity is given by a frequency-dependent kernel (Fig. 4). Frequency is a proxy for depth; the higher
Fig. 1. The geology of Australia at different scales and with different definitions of the units. (a) Most detailed representation considered in this paper. The dashed line is the Tasman Line, which divides the Precambrian western Australia from Phanerozoic eastem Australia (from Zuber et al., 1989). (b) Australian crustal elements, representing continent-scale groups of geophysical domains (from Wellman, 1998). CA, Central Australia; NA, Northern Australia; P: Pinjarra; WA, Westem Australia; SA, South Australia; T, Tasman; NE, New England. (c) A coarse, fourfold regionalization of the Australian continent based on crustal age. The crustal age decreases from Archean in the westernmost part, to predominantly Proterozoic in the central part, and Phanerozoic to the fight of the Tasman line.
22
F.J. Simons et al./Lithos 48 (1999) 17-43
the frequency of the waves the more their sensitivity shifts to shallower depths. For a given frequency, the fundamental mode surface waves (Fig. 4a) sample shallower structure than the higher modes (Fig. 4b), and inclusion of the latter thus provides for increased depth resolution.
3. Partitioned waveform inversion (PWI) The interpretation of waveforms in terms of aspherical variations in Earth's structure is often partitioned into (1) a non-linear waveform inversion that, for each record, seeks to determine the 1D structure
SS SSS
L
/
/
"
.~_
m 36
b
/t
SSS
Fig. 3. Body and surface-wave phases. (a) Composite record section of vertical-component seismograms arranged in order of increasing epicentral distance to two different earthquakes. The dashed lines represent the arrival times of the body-wave phases S, SS and SSS calculated from the ak135 wave speed model (Kennett et al., 1995). The Rayleigh-wave surface wave L R can be thought of as a limiting case of multiple reflections propagating across the surface. The solid lines and shaded areas indicate two group velocity windows. The first window, defined between 4.9 and 4.2 km s-~ selects phases sampling the upper-mantle (the "higher-mode window"). The second window, between 4.2 and 3.4 km s-~ selects fundamental-mode surface waves. We note that for each record the precise windows are set manually upon visual inspection of the waveforms. Seismograms have been filtered between 10 and 45 mHz for the higher modes, and between 10 and 25 mHz for the fundamental modes. (b) Ray geometry in the upper mantle of the body-wave phase shown in (a).
Simons et al./Lithos 48 (1999) 17-43
23
400
400
400
Fig. 4. Sensitivity (Fr~chet) kernels for surface wave propagation. The kernels represent ac/a/3(r), i.e., the sensitivity of the phase velocity c of a particular set of surface wave modes to a perturbation of the shear-wave speed /3 at a particular depth. (Top) Fundamental mode. (Bottom) Higher modes. Surface-wave studies done with fundamental modes are mostly sensitive to shallow structure, while most of the sensitivity at depth is due to the higher modes.
between source and receiver (see Section 3.1) and (2) a geometrical, linear tomographic inversion which combines the individual path constraints into a 3D model for Earth structure (see Section 3.2). Partitioned Waveform Inversion (PWI), developed by Nolet (1990), has previously been used to study the upper mantle beneath Europe (Nolet, 1990; Zielhuis and Nolet, 1994a,b), South Africa (Cichowicz and Green, 1992), North America (van der Lee and Nolet, 1997a,b), and Australia (Zielhuis and van der
Hilst, 1996). Here we review some basic aspects of the technique; for more complete descriptions we refer to Nolet et al. (1986), Nolet (1990), and Zielhuis and Nolet (1994a).
3.1. Waveform inversion for path-averaged structure In the first step of PWI individual seismograms are analyzed and, within the restrictions discussed below, inverted for shear velocity variations with
24
F.J. Simons et al./Lithos 48 (1999) 17-43
depth, 6fl(r) (with r radius), averaged along the source-receiver path. This involves the matching of observed waveforms with theoretical (synthetic) seismograms, which are computed as a sum of surfacewave modes using the JWKB approximation (Woodhouse, 1974). The waveform synthesis requires information about the earthquake focal mechanisms, which is obtained from the Harvard Centroid Moment Tensor (CMT) catalog (Dziewonski et al., 1981) and the National Earthquake Information Center (NEIC) (Sipkin, 1994). The JWKB formalism assumes that lateral heterogeneity is sufficiently smooth compared to the wavelength of the seismic waves used, which imposes a lower limit for frequency. For our application, with perturbation scale lengths typically greater than ~ 400 km, epicentral distances ranging from 1000 to 4000 km, and a typical phase speed of about 4.5 km s-1, a lower frequency limit of not much less than 10 mHz is predicted on theoretical grounds (Kennett, 1995; Wang and Dahlen, 1995; Dahlen and Tromp, 1998). In practice this criterion can be relaxed a bit and for the fundamental modes we consider a lower limit of 5 mHz (although most data are for frequencies higher than 8 mHz). The path-average approximation also implies an upper-frequency limit, both for the fundamental mode and the overtones, because cross-branch coupling between modes, as would arise from lateral heterogeneity (Kennett, 1984; Li and Tanimoto, 1993; Marquering and Snieder, 1995) and which brings out the ray character of the higher modes, such as SS, is not accounted for. The detrimental effects of ignoring mode coupling are aggravated with increasing frequency. However, Marquering et al. (1996) and Marquering and Snieder (1996) have demonstrated that if data coverage is dense the results will be little changed by the adoption of mode coupling techniques in the inversion. Moreover, the largest differences would occur outside the depth range that is also constrained by the horizontally propagating fundamental modes; therefore, we only discuss structure to a depth of 400 km. (We note that structure beneath stations or events that is not well sampled by rays from different directions, such as on the edge of the model, may exhibit spurious vertical structure resulting from the one-dimensionality of the kernels. This may be the case for the deep fast anomaly beneath the NWAO station at Narrogin,
West Australia.) At high frequencies the fundamental-mode Rayleigh waves become sensitive to steep gradients in shallow structure (Fig. 4), such as at the transition from oceanic to continental crust, which can severely distort the waveforms. Along with lowpassing the data, we minimized such effects by accounting for variations in crustal thickness within the region under study on the basis of crustal thickness information from converted phases recorded at the SKIPPY stations (Shibutani et al., 1996; Clitheroe et al., submitted). Kennett (1995) and Kennett and Nolet (1990) conclude that with an upper frequency limit of 20 mHz for the fundamental mode and 50 mHz for the higher modes the surface-mode summation used in PWI provides a representation of the seismic wave field that is adequate for our purposes. Because the admissible frequency limits differ we use group velocity windows to isolate the fundamental and higher mode part of the records so that we can analyze them within the frequency bands discussed above (see also Fig. 3). An example of waveform fits is given in Fig. 5. The wave trains for the fundamental and higher modes have been normalized to unit amplitude. Both sections of the seismogram are fitted separately within the applicable frequency band. The initial fits depict the difference between the observed data (thick solid lines) and the synthetic records (thin lines) produced from a reference model. For this event in Southern Sumatra the fundamental-mode Rayleigh waves arrive earlier than predicted from the average Earth for all stations considered, suggesting relatively fast wave propagation across the western and central part of the continent. But at some stations the mismatch is larger than at others; compare, for instance, the records of ARMA (Armidale, New South Wales) and CTA (or CTAO, Charters Towers, Queensland). These differences are indicative of lateral variations in wave speed (for these earthquakes the data indicate that the average wave speed along the path to ARMA is faster than along the one to CTA). For the same stations, the difference between the observed and synthesized overtones are smaller than for the fundamental mode, suggesting that the wave speed variations in the deep part of the model are smaller than in the shallow part. The bottom panel of Fig. 5 shows the excellent (final) fit of the
25
F.J. Simons et al. ,/Lithos 48 (1999) 17-43 I
I
,
~~~~t.~L~
I
I
I
I
I
I
ar ma
I
Time (s) Fig. 5. Waveform fitting. Observed data are plotted as thick solid lines, predictions are depicted with thin lines. Seismograms are for a Mb 5.8 event in Southern Sumatra, at (103.9~ 5.7~ located at 56 km depth. Great circle paths to the stations are plotted in the inset. "Initial fits" are predictions made by surface-wave summation, using an assumed reference model. "Final fits" are obtained by the nonlinear inversion procedure described in Nolet et al. (1986) (see text).
waveforms, both for the higher-frequency overtones as well as for the lower-frequency fundamental modes.
3.2. Tomographic inversion for 3D structure Once the path-average of the variation of wave speed with depth (6~(r)) is determined for each source-receiver combination these 1D profiles are used as observations in a tomographic inversion for
3D variations in shear wave speed (6~(r,0,q~), with r radius, 0 colatitude, and q~ longitude). This linear tomographic inversion may be performed by a variety of methods, such as the ones described by Nolet (1990), Zielhuis (1992), Zielhuis and Nolet (1994b), or van der Lee and Nolet (1997b). The tomographic problem is parameterized by means of local basis functions (equal-area blocks in latitude and longitude direction; and box-car and triangular basis functions for radius). The system of
26
F.J. Simons et al. / Lithos 48 (1999) 17-43
equations is solved in a generalized least square sense, using the LSQR iterative algorithm (Paige and Saunders, 1982; Nolet, 1985). The damping applied is a combination of first-order gradient damping, which minimizes the differences in structure between adjacent cells, and norm damping, which produces a bias toward the reference model used. The model presented in this paper results from careful experimentation with all variables involved (damping, weights, cell size), and parameters such as variance reduction and retrieval of synthetic input models have been used as guidance. After 200 iterations of the LSQR inversion a variance reduction of about 90% was obtained. The reference model used for the 3D inversion has a crustal thickness of 30 km, which is a reasonable average for the region under study. For the mantle we used a modified PREM model (Dziewonski and Anderson, 1981), smoothly interpolated over the 220 km discontinuity (Zielhuis and Nolet, 1994b).
3.3. Differences with respect to previous studies of Australia We have made several modifications to the methods by Zielhuis and Nolet (1994b), including a different form of weighting of the individual data fits, a parameterization with an increased number of basis functions in radial direction, and we added a parameter that can absorb effects of epicenter mislocation. The effects of these improvements are, however, subtle, and most of the differences with previous models (Zielhuis and van der Hilst, 1996; van der Hilst et al., 1998) can be attributed to the use of an expanded data set. Firstly, we assign an uncertainty to the individual fits based on a weighted combination of the signal bandwidth, the length of the group velocity windows, the xZ-norm, and the zero-lag cross-correlation value of the synthetic and observed waveforms, as well as the ability to fit both the fundamental and higher mode data. The reciprocals of the uncertainties obtained were used to weight the data in the inversion. Secondly, like Zielhuis and Nolet (1994b) we used a combination of boxcar and triangular basis functions, but we have added additional node points in order to extract more information on deep structure from the large number of higher modes in our
data set. Thirdly, we have used the hypocenter locations from the global data file by Engdahl et al. (1998), but in this region source mislocations can be substantial owing to sparse station coverage of the southern hemisphere. A mislocation of 20 km, on an epicentral distance of 2000 km can cause a spurious wave speed anomaly of 1%. Without attempting a formal earthquake relocation we aimed to absorb such effects of source mislocation in a denuisancing parameter in the linear inversion for 3D structure. This has an effect similar to damping, and varying the degree of source relocation allows us to identify the structural features in our model that are required by the waveform data only.
4. Data used in this study Between May 1993 and October 1996 the Australian National University operated the Skippy seismometry project (van der Hilst et al., 1994). This project involved 6 arrays of up to 12 portable broadband seismometers that together synthesized a nationwide array (Fig. 2) and was intended to exploit Australia's location with respect to regional seismicity and provide dense data coverage for a range of tomographic imaging techniques. The individual arrays were deployed for about 6 months at a time. Parts of the large set of surface wave data were used in previous studies (Zielhuis and van der Hilst, 1996; van der Hilst et al., 1998), which focused on central and eastern Australia because data coverage in the west was not satisfactory at the time. In addition to the data from the S~PPY experiment we have used data from broadband permanent stations from the IRIs (Incorporated Research Institutions for Seismology), GEOSCOPE, and AGSO (Australian Geological Survey Organization). The ~ 1600 vertical-component seismograms from about 340 seismic events provide excellent data coverage (Fig. 6). Further improvements in the resolution, in particular of the western part of the continent, are still expected, given the continued monitoring of earthquake activity by Acso. We used the portion of the vertical-component seismogram from the arrival of direct S up to, and including, the arrival of the fundamental mode of the Rayleigh wave. This time window includes multiple
F.J. Simons et al./Lithos 48 (1999) 17-43
_10 ~
_20 ~
-30 ~
_40 ~
27
2000 km, but at 4000 km this is SS, and at 6000 km SSS, etc. (Fig. 3). In recognition of the frequency limitations imposed by the approximations implicit in the method (see above), fundamental mode data were fit in the 5-25 mHz range and the higher modes were modeled within 8 - 5 0 mHz. About 40% of all seismograms contained higher-mode windows for which good fits could be obtained.
5. Results
Here we discuss briefly the spatial resolution and the general aspects of our model, compare our results to variations in shear wave speed as inferred from a global inversion, and describe in detail how the shear wave speed varies within several well-defined tectonic units that constitute the Australian continent.
5.1. Spatial resolution
Fig. 6. (Top) Great-circle paths of the 1596 event-receiver combinations used in this study. (Bottom) Path coverage. Path lengths and variance of the directions of the rays crossing in 2~ 2 ~ cells, two indicators of tomographic quality, are combined and expressed on a relative scale (black indicates well sampled; white indicates no sampling).
body wave reflections at the free surface, such as SS and SSS (Fig. 3). The group velocity windows used for the isolation of the fundamental and higher modes are approximately 3.4-4.2 and 4.0-5.0 km s -1, respectively, but for each individual record the exact limits are set after visual inspection. With this selection, the fundamental mode is truncated before the arrival of scattered and multipathed surface waves, and body waves with turning points in the deep mantle are excluded. For example, direct S is the first body wave phase considered at a distance of
Image quality does not only depend on the sheer number of data but also on the azimuthal distribution of the (crossing) paths (Aki and Richards, 1980; Menke, 1989; Lay and Wallace, 1995). For each cell, we added the variance of the directions (between 0 and ~r) of the rays (normalized from 0 to 1) to the normalized sum of the path lengths to provide a better measure of the quality of data coverage (Fig. 6). The data coverage is good throughout the Australian continent but degrades towards the southwest. Since all AGSO data have not yet been used we expect further improvements in this part of the continent. In order to assess the reliability of the images we have performed test inversions with synthetic data calculated from different input models; the ability to reconstruct an input model from the synthetic data is then used to assess how well real structure can be constrained by the available data. We have used different synthetic input models, with harmonic wave speed variations as well as spike tests (for a discussion of such resolution tests, see, e.g., Humphreys and Clayton, 1988 or Spakman and Nolet, 1988), and others in which we tested the robustness of a specific structure in the model. Based on these tests and on theoretical considerations (Kennett and Nolet, 1990; Zielhuis, 1992; Zielhuis and van der Hilst,
28
F.J. Simons et al. / Lithos 48 (1999) 17-43
1996) we conclude that the horizontal resolution in the best resolved parts of the continent approaches 250 km, and the vertical resolution ranges from 50 km (at 100 km depth) to between 100 and 150 km (around 300 km depth). As expected from the data coverage (Fig. 6), the resolution in the central and east is still superior to that in western Australia. For the purpose of this paper we evaluated whether the data can constrain the wave speed variations on the length scales considered in the finest regionalization described below (see Section 5.4). In the experiment, different wave speeds were assigned to various tectonic regions of the Australian continent. Fig. 7 displays the results of two such tests. In the first, the
input anomalies of Fig. 7c were put at 80 km depth, with zero perturbations elsewhere in the model. We calculated synthetic data for all 1600 paths (both for fundamental modes and the overtones) and repeated the linearized inversion for 3D structure. The result for that layer is given in Fig. 7a. In the second experiment the input pattern (Fig. 7c) was placed at 210 km depth, with the response shown in Fig. 7b. These tests show that the wave speed variations are well resolved on the length scales we are interested in, but at shallow depth the image quality is better than at larger depth. In general wave speed contrasts are well resolved, but the amplitude of the wave speed variations is less well determined. We note
-1 -2 -3 Fig. 7. Results of resolution experiment. Input models were constructed by assigning constant wave speed anomalies to different geotectonic regions and at different depths. Anomalies are in percent from a spherical reference model. Synthetic data were generated on the basis of all wave paths used in actual inversion. (a) Recovery of anomaly placed at 80 km. (b) Recovery at 210 km. (c) Input anomaly used in both tests. Note the difference between both color schemes. Thick dashed lines give approximate location of Tasman Line.
29
F.J. Simons et a l . / Lithos 48 (1999) 17-43
t h a t in t h e r e g i o n a l i z e d (see
Section
presentation
5.4) the wave
speeds
of our results were
averaged
o v e r t h e r e g i o n s i d e n t i f i e d , b u t in t h e t e s t i n v e r s i o n s shown here such averaging was not done.
-1 -2 -3
Fig. 8. Depth slices through our preferred velocity model. Anomalies are in percentage from a spherical reference model (see text). The reference velocities used are 4500 m s -1 at 80, 120, and 170 km depth (a-c); 4513.5 m s -1 at 210 km (d); 4581 m s -1 at 260 km (e), and 4851.5 m s-1 at 400 km depth (e). The color scales used are for a and b, c and d, and e and f, with diminishing values of saturation which reflects the decreasing magnitude of the anomalies with depth. Thick dashed lines give approximate location of Tasman Line.
30
F.J. Simons et a l . / Lithos 48 (1999) 17-43
5.2. Shear waue speed uariations in the Australian upper mantle
Figs. 8 and 9 display some of the tomographic results. The shear wave speed anomalies (/3(r,0,q~) - / 3 0 ( r ) ) are plotted as percentages of the wave speed in the reference model (/30(r)) which is a modified version of PREM (without the discontinuity at 220 km). Wave propagation is slow beneath the Phanerozoic eastern part of Australia and fast beneath the Proterozoic and Archean domains, at least to a depth of 200 km (e.g., Fig. 8a-c and Fig. 9). The Phanerozoic high wave speed lithosphere is
relatively thin, generally less than 80 km (note the cross-sections have been cropped to show the depth range between 15 and 450 km), and overlies a pronounced low-velocity zone that extends to approximately 200 km depth (Fig. 9) (see also Goncz and Cleary, 1976; Goncz et al., 1975; and Zielhuis and van der Hilst, 1996). In agreement with Zielhuis and van der Hilst (1996) our model reveals that at depths shallower than 150 km the wave speed gradient from the eastern to the central domains occurs east of where it would be expected on the basis of the exposure at the surface of Proterozoic rock (indicated by the
o
ig)
Fig. 9. Profiles through the model. Anomalies in percentage from a spherical reference model (see text). The high wave speeds in the northeastern corner of the map, which show at the fight hand side of cross-sections (a) and (c), reflect the recent subduction of the Pacific beneath the Indo-Australian plate. The high wave speed feature to the southwest of Tasmania, (c), may be related to the Australia-Antarctic discordance (Gurnis et al., 1998) but this part of our model is not well sampled (see Fig. 6).
F.J. Simons et al./Lithos 48 (1999) 17-43
Tasman Line). The presence of high wave speed lithosphere east of the Tasman line was confirmed by the analysis of fundamental mode dispersion between stations along the same great-circle path (Passier et al., 1997) (Fig. 10). In several regions, the lateral wave speed contrast coincides with surface outlines of sedimentary basins in easternmost Australia, in particular the western margin of the Bowen and Surat basins. At greater depths the wave speed divide shifts westward, and at 170 and 210 km depth it appears to parallel the Tasman line and the western margin of the Eromanga basin. In northeastern Australia, the wave speed contrast is located to the west of the Georgetown Inlier, and this Proterozoic unit is
31
part of a rather thin high wave speed lid with a pronounced low velocity zone underneath (see Fig. 9). In contrast, the Proterozoic shields of central Australia are delineated by high wave speeds down to at least 250 km (e.g. between stations WRAB and SCOG in Fig. 10). The area of the Late-Paleozoic Alice Springs orogeny (Amadeus basin and Musgrave and Arunta blocks) stands out from the adjacent shields by lower wave speeds. This may reflect the thick layer of sediments in this region. Interestingly, also the Kimberley block, which is often interpreted as the westward continuation of the North Australia craton (see, e.g., Shaw et al., 1995 and Fig. l b) has a seismic signature that differs significantly
Fig. 10. Stations, events, great circle paths, and wave speed profiles at selected locations, superimposed on the shear wave speed variations at a depth of 140 km depth (Zielhuis and van der Hilst, 1996). The panels display wave speed variation to a depth of 250 km in eastern and central Australia as inferred from the differential dispersion of fundamental model Rayleigh waves between stations along the same great circle path (Passier et al., 1997). In all velocity panels, perturbations range from - 5 % to 5% and the depth ranges from 0 to 250 km. (Modified after Passier et al., 1997). Note that our current models are based on more data than were available to Zielhuis and van der Hilst (1996).
32
F.J. Simons et a l . / Lithos 48 (1999) 17-43
from that of the Proterozoic craton. Likewise, the deep structure beneath the Canning basin differs from that b e n e a t h central Australia and from the A r c h e a n cratons further South (Fig. 8 b - d ) . B e l o w 200 k m depth the central cratons continue to be m a r k e d by fast anomalies, but the A r c h e a n Yilgarn, Pilbara, and G a w l e r cratons do not seem to be m a r k e d by wave speeds that are significantly higher than the reference model, with the exception of the localized high wave speeds under station
N W A O in the southern Yilgarn craton, but these might be an artifact of using o n e - d i m e n s i o n a l sensitivity kernels (see Section 3.1). W i t h the exception of the region e n c o m p a s s i n g the Alice Springs orogeny, the mantle beneath 300 k m depth beneath the Proterozoic and A r c h e a n appears to be rather h o m o g e n e o u s . In this depth range, the most pron o u n c e d fast anomalies are associated with the subduction zones to the North and Northeast of Australia and with the intriguing structure east of the Mt.
-5* _10 ~ _15 ~ _20 ~ _25 ~ _30 ~ -35 ~ _40 ~ -45 ~ -2
9 1
!
e'~--4 k..
~-5 2 _o
..............................
-2
-4
-6 W a v e l e n g t h (km)
Fig. 11. Comparison with global model in spatial domain. Clockwise: (a) Rayleigh-wave inversion result by Ekstr~im et al. (1997), at 60 s. This phase-velocity map records the anomalies of fundamental modes of some (narrow) frequency band around 60 s, or 17 mHz. From inspection of the kernels of Fig. 4, such map represents structural wave speed deviations with a maximum sensitivity around 120 km depth. We therefore compare it to our inversion results at those depths. (b) Wave speed anomalies at 120 km, using only fundamental modes from permanent seismometer stations (from IRIS and GEOSCOPE).The resolution of the result is roughly equivalent to the one in (a), with certain differences attributable to differences in path geometry. In (a) and (b), the location of the 6 permanent stations is indicated by the black triangles. (c) Our preferred model at 120 km depth. Thick dashed lines give approximate location of Tasman Line. (d) Power spectrum of the shear wave speed model at 80 km depth. The results are compared with the model by EkstriSm et al. (1997).
33
F.J. Simons et al./Lithos 48 (1999) 17-43
Isa block (at 23~ Resolution tests and inversions with subsets of the data demonstrate that the latter anomaly is well resolved and required by the data used. 5.3. G l o b a l vs. regional m o d e l s
Our regional model reveals variations in shear wave speed on smaller scales and with larger amplitude than indicated by global tomography models. To better understand such differences we compared our maps to the global images both in the spatial and spectral domains. Fig. 1 l a displays the wave speed variations according to the global model by EkstriSm et al. (1997), which contains information from about
1700, globally distributed earthquakes, and only a small number of permanent stations on or near the Australian continent. We obtain a similar result if we apply our inversion technique to fundamental mode data from the six regional IRIS and GEOSCOPE stations that were also used by EkstriSm et al. (1997) (Fig. l lb), although we retrieve somewhat lower wave speeds beneath the marginal basins east of Australia. Following Chevrot et al. (1998a,b) we also compared the power spectrum of our model with that from EkstriSm, which was parameterized in spherical harmonics up to degree 40, corresponding to a wavelength of about 1000 km. Across the entire spectrum our regional model does have somewhat higher amplitudes than the global model by EkstriSm et al.
-1 -2 4
-1 _45 ~
1170 km ,
-2 130 ~
140 ~
150 ~
160 ~
120 ~
130 ~
140 ~
150 ~
Fig. 12. Four-part regionalized representation of the shear wave speed model. This regionalization is based on large-scale variations in crustal age (see Fig. l c).
34
F.J. Simons et al./ Lithos 48 (1999) 17-43
(1997), but particularly so at wavelengths smaller than 2000 km (Fig. 1 l d). This is primarily a consequence of degrading resolution at those wavelengths in the global model (see Fig. 13 in EkstriSm et al., 1997). From this comparison we conclude that when evaluated at similar wavelengths the spatial and spectral characteristics of global and our regional models are generally in good agreement with each other, and the superior resolution at high wave numbers (comparing, e.g., Fig. 1 l a and c) is not dictated by the parameterization or damping but results from the dense data coverage provided by the Skippy
project and the inclusion of higher-mode Rayleigh waves. 5.4. Seismic signature o f distinct tectonic regions
" T e c t o n i c " regionalizations, whereby seismic properties are averaged over well-defined geological domains, have formed the basis of first-order testing of the tectosphere hypothesis (Jordan, 1975a, 1981b; Okal, 1977; Polet and Anderson, 1995), the synthesis of global seismic data (Nolet et al., 1994), and the calculation of travel-time corrections (Gudmundsson
:..:'.:"":::'"::::::::'"
-
" ".
it
'i., !
_
':%..... .....%....:..::.1 "'; ~
ii ~
;,.
, /
's
s
-"..-..~...".. ~."
i
i ii!:
E 300-
' iii a 400I - :i~i
,:j -
.:?
li li
............
I I_
,,,-,
i| II
-I
4 Fig. 13. Age-dependent wave speed variations with depth. (Left) Absolute wave speeds. (Right) Wave speed perturbations from a continental reference model. (Inset) Geographic location of the regions used (see Fig. lc). Compare with Fig. 12 for a planview. Below ~ 250 km, the differences of average wave speed between those regions are an order of magnitude smaller than in the shallow mantle.
35
F.J. Simons et al./Lithos 48 (1999) 17-43
and S a m b r i d g e , 1998). W e h a v e used regionalizations to investigate the relationship b e t w e e n the
thickness of the high velocity lid and crustal age at different length scales.
-2
-4 4
-2 _40 ~
Fig. 14. Detailed regionalized representation of the shear wave speed model. In (a) and (c), a 12-part subdivision was chosen, whereas in (c-f), the domain centered over the Eromanga basin was subdivided to bring out the anomalous character of the fast structure centered at 23~ For more discussion, see text. Thick dashed lines give approximate location of Tasman Line.
36
F.J. Simons et a l . / Lithos 48 (1999) 17-43
In our first experiment we used a three-fold tectonic regionalization based roughly on crustal age (Fig. l c), with a further subdivision of the eastem domain in two parts motivated by wave speed maps (see Zielhuis and van der Hilst (1996) and Fig. 8). We averaged the wave speed perturbations over these domains. Fig. 12a-b shows the rather dramatic westward progression of seismic wave speed, which correlates well with the general increase in crustal age (Fig. l c). However, deeper in the upper mantle this relationship breaks down and beneath 200 km depth the central Proterozoic is characterized by fast wave propagation while the average wave speeds in the Archean units are comparable to those under the Phanerozoic at large (Fig. 12c,d). Radial wave speed profiles, calculated for each domain, further illustrate that the recently tectonized eastem seaboard of Australia is generally slow, with a pronounced lowvelocity zone centered around 150 km (Fig. 13). The Precambrian central and western parts do not show this reduction in wave speed; instead, the high-velocity lithosphere beneath these provinces is generally 200-250 km thick. Note that the Archean is faster than the Proterozoic for z < 150 km but that the high wave speed signature does not extend to as large a depth (Fig. 13). Our results are in general agreement with the results by Gaherty and Jordan (1996) and Gaherty et al. (1996) who argued for a difference of similar magnitude between fast continental and slow oceanic mantle to a depth exceeding 250 km. It is obvious from geological maps and results of tomography that significant variations in crustal structure and seismic properties (and our ability to constrain them) occur on much smaller length scales than represented by the large scale regions discussed above. Therefore, we designed another regionalization based on geological units that better represent variations in tectonic age. This tectonic subdivision (Fig. 14) is based on the crustal elements map by Shaw et al. (1995) and refinements by Wellman (1998); it consists of a northem, central, westem, and southern Australian domain, as well as the Tasman group and the New England fold belt (see Fig. l b), and the Pinjarra (the southwestern-most part). On the basis of our wave speed maps (Fig. 8) we treat the Kimberley block separately from the northern Australian domain, the Canning basin apart from Central Australia, and the Pilbara apart from the
Yilgam craton. In Fig. 14a,b we maintain our earlier division of the Tasman domain into a unit comprising the (westem parts of the) Eromanga and Murray basins and a unit with the eastem Lachland and New England fold belts (see Fig. l a). In Fig. 14c-f we have subdivided the Eastem and Westem Tasman domain in a northem and southem part so as to prevent the anomalous fast structure centered around (145~ 23~ from biasing a larger region. This regionalization (Fig. 14) largely confirms the westward progression of seismic signature at depths less than 200 km but also reveals significant departures of the general trends. The high wave speed in the Archean western Australian domain can largely be attributed to the Yilgarn block; the Pilbara block is not marked by higher-than-average wave speeds but resolution in this particular area may be lacking. Recall that the high wave speeds beneath the station NWAO may be an artifact of using 1D kemels, and the signature of the Yilgam region may thus be biased. The high wave speed under central Australia is mostly associated with the central block-and-basin structure and the North Australia craton. Beneath the Kimberley block and the Canning basin the pronounced positive wave speed anomalies vanish beneath 150 km depth, and at a slightly larger depth also the Archean Yilgam, Pilbara, and Gawler units lose their deep seismic expression. However, the Proterozoic North Australia craton is still faster than average at depths in excess of 250 km. Surprisingly, the wave speed anomaly beneath the region encompassed by the Alice Springs orogeny seems to persist to depths in excess of 300 kin. The upper mantle and transition zone beneath Eastern Australia seems more complex than its Precambrian counterpart (see also Fig. 9c and Section 5.2) and on average the transition beneath geological domains in eastern Australia are faster than in the west (Fig. 9b,c).
6. Discussion
The tectosphere concept as formulated by Jordan (1988) implies that the negative thermal buoyancy of the CLM is largely offset by changes in composition. Here we do not image composition and, therefore, use a less restricted definition. High wave speed lids that are more than ~ 200 km thick would be dynam-
F.J. Simons et al. ,/Lithos 48 (1999) 17-43
ically unstable if the negative buoyancy owing to the low temperature is not somehow compensated by changes in composition (Jordan, 1975a,b; Anderson and Bass, 1984; Polet and Anderson, 1995; De Smet et al., 1999, this volume; Shapiro et al., 1999a,b, this volume). In the following we loosely refer to "continental keels" or "tectosphere" when the thickness of the high wave speed lid as inferred from our results exceeds this value.
6.1. Regional deviations from a global pattern On the basis of global wave speed models, Jordan (1981b), Nolet et al. (1994), Polet and Anderson (1995), and others, concluded that there is a good agreement between the velocity anomalies and the surface pattern of crustal formation ages. The Precambrian domains are fast, whereas most Phanerozoic regions are underlain by a seismically slow upper mantle. Such observations have often been used to corroborate the correlation between crustal age and the composition of the CLM inferred from geological data. Global wave speed models have suggested that the high wave speeds persist down to 500 km depth (Su et al., 1994; Masters et al., 1996) but the depth extent of the high-velocity anomalies under stable continental regions has not been well constrained. The fundamental and higher mode data used in our study are not consistent with these earlier results and, instead, constrain the thickness of the high wave speed CLM to less than 250-300 kin. For Australia a strong correlation between surface age and seismic structure does exist, but only at long wavelengths (A > 1500 km) and for depths shallower than ~ 200 km. At smaller length scales there are parts of Precambrian shields that are marked by high wave speeds to even larger depth but there are also many such regions without a deep seismic expression. At the length scales considered in our study, the "tectosphere" beneath the Proterozoic domains is as thick as or thicker than beneath Archean units, but the wave speed in the latter appears to be higher (Fig. 13). However, some Proterozoic domains, for instance the Georgetown Inlier in northern Queensland, are not marked by a deep, seismologically fast keel. On the other hand, some regions devoid of Proterozoic or older outcrop appear to have a tecto-
37
sphere-like CLM. Significant variations in seismic properties thus occur beneath geological provinces of comparable crustal age or tectonic setting and similar complexity may exist for other continents as well. Hence, for Australia there does not seem to be a simple relationship between tectosphere structure and crustal age.
6.2. Eastern Australia Beneath easternmost Australia the high wave speed lid is thinner than 80 km in most places, entirely absent in some (for instance beneath the Queensland Volcanic Province), and mostly underlain by a pronounced low velocity zone centered at about 150 km depth (see Zielhuis and van der Hilst (1996) and references therein) (see also Figs. 9 and 13). Tectonically this area is relatively young, and in the southeast volcanism may have ceased only a few thousand years ago (Johnson, 1989). The geothermal gradient is steep, the present-day surface heat-flow is high (70-100 mW m - 2 ) (Cull and Denham, 1979), and studies of mantle conductivity (Lilley et al., 1981; Finlayson, 1982) suggest that volatile-rich material or partial melt may be present. Therefore, the pervasive lower-than-average wave speeds in this region are most likely the result of thermal processes associated with the Late-Cretaceous ( ~ 80 Ma) opening of the Tasman Sea and accompanying volcanism (Storey, 1995), volatile infusion during subduction prior to that event, and the recent volcanism in the region. To the west of the low wave speed region, but still east of the Tasman line, the wave speed signature suggests the presence of a seismically fast root. Although it is not as thick as the CLM beneath the Proterozoic proper, this could imply that some of the early Paleozoic fold belts (for example the Adelaidean and western part of the Lachland) and basins (e.g., western part of the Murray basin) are underlain by Precambrian basement (Zielhuis and van der Hilst, 1996). This interpretation, which was recently corroborated by Re-Os model ages of up to 1.96 Ga for upper mantle xenoliths in western Victoria (Handler et al., 1997), would have implications not only for understanding the structural development of these fold belts and basins but also for the paleogeographic reconstruction of the Rodinia supercontinent (Powell, 1998; K. Karlstrom, pers. comm., 1999).
38
F.J. Simons et al./Lithos 48 (1999) 17-43
It is unlikely that the anomalously fast structure at a depth greater than 250 km, i.e., beneath the pronounced low velocity zone (see Fig. 8c-fFig. 9b), is related to stable lithospheric structure formed in situ. Instead it may represent allochthonous material. Perhaps it represents one or more accreted terranes of continental tectosphere that was originally formed elsewhere. Given the north-northeastward motion of the Australian plate relative to the lower mantle (DeMets et al., 1990), a perhaps more plausible explanation is that fragments of formerly oceanic lithosphere that were subducted beneath the island arcs to the north and northeast of Australia have subsequently been overridden by the continent. 6.3. Central and Western Australia
As mentioned above, the seismic signature of the tectosphere in central and western Australia changes significantly within geological provinces of comparable age and tectonic setting, and the thickest tectosphere appears to be associated with the Proterozoic shields, not with the Archean cratons. Moreover, some of the fast anomalies beneath central and eastern Australia do not correlate with surface geology in an obvious fashion. A detailed interpretation of these observations is beyond the scope of this paper, but we speculate on a few scenarios that might explain the observed complexity. Assuming that the interpretation of the thick high wave speed structure in terms of stable continental keels (tectosphere) is basically correct, two fundamentally different explanations can be given. One possibility is that the differences within and between the Archean-Proterozoic provinces are entirely due to how the structures were originally formed, and have persisted ever since. Conditions for melt depletion associated with tectosphere formation could have varied from region to region. However, this complexity is not readily consistent with the correlation between crustal age and mantle composition inferred from geological data (e.g., Griffin et al. (1998)), which implies a certain uniformity of formation processes. Another explanation is that a thick (z > 250 km) tectosphere formed everywhere but only survived beneath central Australia. Shapiro et al. (1999a) showed that a cold tectosphere could be supported by high viscosity alone but that compositional buoy-
ancy is required to annihilate part of the gravity signal and to stabilize the "keel" during formation when the temperature was higher and the viscosity lower. The CLM can be eroded by later thermal processes, which would reduce its viscosity, in particular when it is already weakened by hydratation, for example by volatile infusion during subduction (Nolet and Zielhuis, 1994). The removal of CLM by plumes, subduction, or small-scale convection associated with continental break-up has been invoked to explain the deep structure beneath western part of the Russian Platform (Nolet and Zielhuis, 1994), the anomalously low subcrustal Pn and Sn velocities beneath the southwestern margin of the Proterozoic Baltic Shield (Bannister et al., 1991), and the absence of a seismically fast CLM beneath the Rocky Mountain front in North America (Eggler et al., 1988), the Proterozoic Grenville province in the eastern part of United States (van der Lee and Nolet, 1997b), and the Archean Sino-Korean craton (see Griffin et al., 1998 and references therein). The position of the "rootless" Archean Yilgarn, Pilbara, and Gawler blocks near passive margins suggests that an original deep high-velocity lid could have been destroyed by convective processes associated with continental break-up when western and southern Australia parted from India and Gondwana (onset ~ 130 Ma) and Antarctica (starting ~ 80 Ma) (Veevers, 1984). Much like the rifting of Australia from New Zealand, the separation of Greater India from Australia was preceded and accompanied by extensive magmatism and related thermal processes (Storey, 1995). Alternatively, the deep (z > 200 km) high wave speeds represent temperature perturbations associated with subcontinental downwellings instead of stable tectosphere. This may apply to the localized high wave speed structure beneath the central Australian region involved in the Late-Paleozoic Alice Springs orogeny, which may continue to lower mantle depths (Grand, 1998, pers. comm.). In this scenario, the permanent thick part of the continental plate might have an almost uniform thickness of about 200-250 km beneath all domains of Archean or Proterozoic age, and the deep, high-velocity anomalies under central and northern Australia reflect convective instabilities under the continent. Cold convective drips beneath continents have been sug-
F.J. Simons et al./Lithos 48 (1999) 17-43
gested before (Gurnis, 1988; De Smet et al., 1999, this volume) and were invoked to explain the deep high wave speeds beneath the Canadian shield (Pari and Peltier, 1996) (see, however, Shapiro et al., 1999b). For North America, Li et al. (1998) showed that the undulations of the 410 km discontinuity, which marks the isochemical phase change from olivine to wadsleyite, are small and do not correlate with the presence of a cold CLM. From this they conclude that the cold drips must be small, if they exist at all. For Australia such studies have not yet been completed.
7. Conclusions
We have analyzed about 1600 broad-band, vertical-component seismograms provided by the permanent IRIS, GEOSCOPE, and AGSO stations and the portable seismograph stations of the SKIPPY project. The combination of dense wave-path coverage and the use of both the fundamental and higher modes of the Rayleigh (surface) wave allowed us to delineate aspherical wave speed variations of the Australian continent in unprecedented detail. Our current inversions account for isotropic wave speed variations only. In this paper we have focused on the presentation of the results, with special emphasis on the relation between wave speed variations in the upper mantle and the geological age of the overlying crust. In order to investigate whether that relationship depends on length scale of the wave speed variations and crustal structures considered, we have performed different tectonic regionalizations of our new wave speed model. We reach the following conclusions: (1) For large scale lengths, and above ~ 200 km, there is a dramatic westward progression of lithospheric wave speed with increasing crustal age, in agreement with global studies. In general, the Phanerozoic lithosphere ( < 80 km thick) is underlain by low velocity anomalies (with respect to a continental reference model); such a LVZ is absent beneath the Precambrian units, which instead are underlain by fast S-wave anomalies to about 200-300 km depth. (2) Even on these large scales (A > 1500 km) the high wave speed lid of the Proterozoic provinces (central Australia) is generally thicker (250-300 km)
39
than that of the Archean cratons (175-250 km), but the wave speed anomalies are higher in the latter. Our estimates of the thickness of the Archean CLM are slightly larger than, but given the uncertainty, consistent with, those of the Kaapvaal craton (Priestley, 1999, this volume). (3) The Early-Paleozoic mountain belts in southeastern Australia (Adelaidean and western Lachland) and the western part of the Phanerozoic Murray basin are underlain by tectosphere-like upper mantle ( ~ 150 km thick). The inference that crustal deformation has occurred on a Proterozoic ramp is consistent with mid-Proterozoic Re-Os ages for upper mantle xenoliths from the Mt. Gambier volcanic complex several hundred km east of the outcrop divide (the so-called Tasman line) between the Proterozoic central shields and the Phanerozoic fold belts (Handler et al., 1997). These observations are important for understanding the extent of Precambrian lithosphere in Australia as well as the reconstruction of the Proterozoic super-continent Rodinia. (4) The high wave speed structures that we interpret as old continental keels often continue off-shore, for instance northward into New Guinea, and southward into the Indian Ocean. The latter may represent stretched continental lithosphere resulting from rifting between Australia and Antarctica. Likewise, the inferred high wave speed keel just east of the Tasman Line could represent thinned lithosphere due to extension associated with the break-up of Rodinia (post-750 Ma). (5) At shorter wavelengths, parts of the Precambrian shields are underlain by high wave speeds to depths greater than 200 km but other such domains lack a deep seismic signature altogether. (6) The large variability of the seismic signature of the deep continent may be explained by (1) as yet poorly understood differences in the original fractionation and depletion of the upper mantle and m more likely (2) the local disruption and removal of once thick tectosphere by later tectonic processes. We notice that the "rootless" Archean cratons are located near old passive plate margins, suggesting that destabilization of tectospheric mantle has occurred by small-scale convection associated with continental break-up and subsequent rifting. Alternatively, some of the deep high wave speed structures may reflect cold downwellings (e.g., beneath central
40
F.J. Simons et al./Lithos 48 (1999) 17-43
Australia) or the relics of allochthonous terranes (either exotic tectospheric blocks accreted onto eastem Australia, or remnants of subducted slab overridden by the plate on its northeastwardly joumey). (7) Variations in deep continental structure within regions of similar crustal age and tectonic setting are significant and often as large as variations between such age units. This implies that the wave speed models do not always support a simple relationship between mantle composition and crustal age and that global correlations based on geological regionalizations on scales smaller than about 1000 km should be considered with caution.
Acknowledgements We are grateful to Brian Kennett at the Australian National University for his support of the SKIPPY project and for many discussions about the structure of the Australian continent. We thank Mark Leonard at the Australian Geological Survey Organisation for help in making the AGso-data available to us. We thank Erik Larson, Roberta Rudnick, and Bill McDonough for thoughtful comments on the manuscript, and Bill McDonough for expediting the review process. This research was supported by the National Science Foundation, grant EAR-9614341.
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Zielhuis, A., 1992. S-wave velocity below Europe from delay-time and waveform inversions. PhD thesis, Utrecht Univ., Utrecht, The Netherlands. Zielhuis, A., Nolet, G., 1994a. Shear-wave velocity variations in the upper mantle beneath central Europe. Geophys. J. Int. 117, 695-715. Zielhuis, A., Nolet, G., 1994b. Deep seismic expression of an ancient plate boundary in Europe. Science 265, 79-81. Zielhuis, A., van der Hilst, R.D., 1996. Upper-mantle shear velocity beneath eastern Australia from inversion of waveforms from S ~ P Y portable arrays. Geophys. J. Int. 127, 1-16. Zuber, M.T., Bechtel, T.D., Forsyth, D.W., 1989. Effective elastic thicknesses of the lithosphere and mechanisms of isostatic compensation in Australia. J. Geophys. Res. 94, 9353-9367.
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LITHOS ELSEVIER
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Lithos 48 (1999) 45-56
Velocity structure of the continental upper mantle" evidence from southern Africa K. Priestley Department of Earth Sciences, Bullard Laboratories, University of Cambridge, Cambridge CB3 0EZ, UK
Received 14 April 1998; received in revised form 20 November 1998" accepted 15 December 1998
Abstract The velocity model for southern Africa of Qiu et al. [Qiu, x., Priestley, K., McKenzie, D., 1996. Average lithospheric structure of southern Africa. Geophys. J. Int. 127, 563-587] is revised so as to satisfy both the regional seismic waveform data and the fundamental mode Rayleigh wave phase velocity data for the region. The revised S-wave model is similar to the original model of Qiu et al. except that the high velocity, upper mantle lid extends to 160 km depth in the revised model rather than to 120 km in the original model. Sensitivity tests of the regional seismic data show that the minimum velocity in the S-wave low velocity zone can be as high as 4.45 km s-1 compared to 4.32 km s-1 in the Qiu et al. model. The vertical S-wave travel time for the revised south African model is compared with the vertical S-wave travel times for the global tomographic models S12WM13 and S16B30, and they are found to be similar. 9 1999 Elsevier Science B.V. All rights reserved. Keywords: Velocity structure; Continental upper mantle; Southern Africa
1. Introduction The southern African region is particularly important in understanding upper mantle properties because of the large number of kimberlite eruptions, primarily on the Kaapvaal craton, which have brought up nodules from upper mantle depths. Such rocks are the only direct information available on the composition of the lower lithosphere; hence it is important to determine whether these upper mantle samples are derived from " t y p i c a l " continental mantle or are unique to the kimberlite source region below the craton. Qiu et al. (1996) [hereafter referred to as QPM96] published an average compressional and shear wave
velocity model for southern Africa. The crust of their model was constrained by published seismic refraction results and receiver function analyses. The upper mantle of the model was constrained by inversion of multi-mode waveforms using the differential locked-mode seismogram technique of Gomberg and Masters (1988) and with further forward modeling of the waveforms using reflectivity (Fuchs and Muller, 1971). Seven of the three-component waveforms analyzed were taken from the long period channel of the SLR D W W S S N station (Fig. 1); the remaining waveforms analyzed were taken from the broadband channel of the SUR, BOSA, LBTB and LSZ seismographs and were lowpass filtered at 0.05 Hz with a 3-pole butterworth filter. The average velocity model
0024-4937/99/$ - see front matter 9 1999 Elsevier Science B.V. All fights reserved. PII: S0024-4937(99)00021-3
46
K. Priestley /Lithos 48 (1999) 45-56
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Fig. 1. Source-receiver paths for regional earthquake seismograms used in this study superimposed on the major crustal subdivisions of southern Africa. Triangles denote seismograph stations; stars denote earthquake locations; and dotted lines denote paths for events used in the regional waveform modeling. Events 1-8 were used by QPM96; event 9 is the September 21, 1997 earthquake (7.25~ 30.16~ 30 km deep). The phase velocity for fundamental mode Rayleigh waves has been measured for the SUR-BOSA and BOSA-LBTB paths.
of the crust and upper mantle derived from the analysis of waveforms from 12 paths largely confined to the Kaapvaal and Zimbabwe cratons and the Limpopo belt (Fig. 1) has the following features. Below the 42 km thick crust there is an 80 km thick, high S-wave velocity upper mantle lid. The P and S
velocities are 8.09 and 4.62 km s -1, respectively, beneath the Moho and the compressional and shear velocity gradients in the lid are 0.0008 s -1 and 0.0013 s -~, respectively. The high velocity upper mantle lid is required to explain both the high velocity observed in southern Africa and the propagation
K. Priestley / Lithos 48 (1999) 45-56
of long period fundamental mode Love and Rayleigh waves from the regional events observed on the paths shown in Fig. 1. Below the lid, QPM96 found a shear wave LVZ in which the S-wave velocity drops to 4.32 km s-1 at 250 km depth. The S-wave LVZ is required to match the arrival time and amplitude of the higher mode waveforms. Forward modeling of Pnl waveforms using reflectivity synthetics shows that an upper mantle P-wave LVZ is not required by the data. Below 125 km depth the P-wave gradient increases to 0.0015 s-1 and increases again to 0.0035 s -~ between 250 km depth and the 410 discontinuity. QPM96 compared the seismic velocity of their upper mantle model with velocities they determined from the analysis of peridotite nodules from the kimberlite volcanics on the Kaapvaal craton. The velocities determined from the nodules coming from the upper mantle lid depth agree well with the seismic structure, but velocities determined from the nodules coming from deeper depths are higher than the seismic velocity at those same depths. The agreement in the lid velocity determined from seismology and petrology suggests that the nodules are representative of the upper mantle lid composition beneath the wider region of southern Africa and are not merely characteristic of the kimberlite source region. Two issues are addressed here. In Section 2, the south African upper mantle velocity model of QPM96 is modified. QPM96 noted that the Rayleigh wave phase velocity curve for their model underestimates the observed dispersion curve of B loch and Hales (1968) for southern Africa. New Rayleigh wave phase velocity measurements are presented which give results similar to those of B loch and Hales (1968) and the QPM96 shear wave model is modified so as to agree with both the regional waveform data and the Rayleigh wave phase velocity data. Section 3 deals with the apparent differences in upper mantle structure between the south African model of QPM96 and those of the global tomographic models. The tomographic models ( S u e t al., 1995; Masters et al., 1996) show higher than average S-wave velocities to depths as great as 500 km in a wide region beneath southern Africa. This differs from the QPM96 model and other upper mantle models for southern Africa derived from regionally recorded seismic data (Bloch et al., 1969; Cichowicz
47
and Green, 1992) which show high S-wave velocities extending only to 120 to 160 km depth. The vertical S-wave travel times through the upper mantle of the QPM96 model and the global tomographic models are compared and found to be similar.
2. Seismic constraints on the southern Africa upper mantle
Bloch and Hales (1968) measured interstation fundamental mode Rayleigh wave phase velocities across the Kaapvaal and Zimbabwe cratons; these measurements, along with fundamental and higher mode Rayleigh and Love wave group velocities, were interpreted by Bloch et al. (1969). Their S-wave models derived by a combination of inversion and forward modeling include a high S-wave velocity upper mantle lid overlying a LVZ; the sub-Moho shear wave velocity is about 4.70 km s -1 , the lid has a positive shear wave velocity gradient, and the base of the upper mantle lid in their model is at 120 km depth. QPM96 used this mantle shear wave model as the starting mantle model for the regional waveform inversion and their resulting velocity model is similar to that of B loch et al. (1969). QPM96 also used one of the Cichowicz and Green (1992) South Africa S-wave velocity models as a starting model for the waveform inversion so as to examine the influence of the initial model on the inversion results. Using this starting model, QPM96 obtained an inversion result similar to that obtained when the B loch et al. (1969) model was used as an initial model, suggesting that the QPM96 inversion model was not strongly dependent on the Bloch et al. (1969) starting model. However, the QPM96 velocity model underestimates the Bloch and Hales (1968) Rayleigh wave phase velocity curve. QPM96 noted this discrepancy, showed that synthetics seismograms computed for the Bloch et al. (1969) model did not match the observed seismograms, and concluded that the inversion of the multi-mode waveform data coupled with the observed S n velocity provided a more stringent constraint on the S-wave velocity structure than did the fitting of the dispersion data on its own. Fundamental mode Rayleigh wave phase velocities have been measured from two events for the LBTB-BOSA path (A 400 km) and two events for
48
K. Priestley / Lithos 48 (1999) 45-56
the BOSA-SUR path (A 597 km) (Fig. 1) using the transfer function method of Gomberg et al. (1988). The upper and lower bounds for these dispersion data, the average dispersion curve of B loch and Hales (1968), and the theoretical dispersion curve for the QPM96 model are shown in Fig. 2. The B loch and Hales (1968) curve lies within the bounds of the new measurements; the QPM96 dispersion curve underestimates both the Bloch and Hales (1968) phase velocity curve and lies below the lower bounds of the dispersion measured here for periods 50-90 s. This suggests that the S-wave velocity in the 120-170 km depth range of the southern Africa model of QPM96 is too low. QPM96 conducted a number of tests to determine the sensitivity of their data to the features of the upper mantle velocity structure. From the multi-mode waveform inversions the average depth to the base of the lid was found to be 120 km. However, their
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Fig. 2. Comparison of observed fundamental and higher mode Rayleigh wave dispersion curves and the theoretical curves for the QPM96 and revised southern African velocity models. The upper and lower bounds of the fundamental mode Rayleigh wave phase velocity measured for the paths SUR-BOSA and BOSA-LBTB are shown as the two thick solid lines. The observed dispersion curves of Bloch and Hales (1968) and Bloch et al. (1969) are denoted by the thin solid lines. The theoretical phase velocity curve for the QPM96 model (dot-dash line) lies below the Bloch and Hales (1968) phase velocity curve and below the lower bound of the phase velocity measured in this study. Increasing the depth to the top of the S-wave LVZ to 160 km results in the theoretical dispersion curves (dashed lines) which match the observed fundamental mode phase velocity curves and the fundamental and higher mode group velocity curves.
sensitivity tests indicate that the upper mantle lid could persist to a depth of ~ 150 km without significantly degrading the fit of the synthetic seismograms to the regional seismic waveforms. A more extensive test of the sensitivity of the data to the depth of the base of the lid is shown in Fig. 3a-c. These tests compare synthetic seismograms for four lid models with the observed seismograms at three distance ranges. Increasing the depth to the top of the S-wave LVZ to 160 km results in a small additional increase in the higher mode amplitude at 1038 km distance (Fig. 33). There is only a small additional increase in amplitude if the depth to the LVZ is increased. None of these lid models produce a significant change in the travel time at this distance. The S-wave at this distance range consists of energy turning in the lid, but the distance range is not great enough to be sensitive to the top of the LVZ. The comparison at 1480 km distance (Fig. 3b) shows that increasing the depth to the top of the S-wave LVZ advances the travel time and increases the amplitude of the synthetic waveform with respect to the observed waveform. Increasing the top of the LVZ to 160 km results in a small travel time advance of the synthetic with respect to the observed waveform. The match is still acceptable but not as good as that of the original QPM96 model. Increasing the top of the LVZ depth to greater depths results in both an earlier and larger amplitude synthetic with respect to the observed waveform. At this distance range most of the S-wave consists of energy turning in the lower lid, and increasing the lid thickness increases both the arrival time and the amplitude of the S-wave. The comparison at 2106 km distance (Fig. 3c) shows an acceptable match of the synthetic and observed waveform for a 160 km deep LVZ, but thicker lids advance the arrival time of the waveform although the waveform shape is not altered. At this distance range the S-wave consists of energy turning in the vicinity of the 410-km discontinuity. Increasing the lid thickness causes an advance in the arrival time but no change in the amplitude. QPM96 also comment that the S-wave LVZ beneath southern Africa was necessary to match the seismograms but that the velocity structure of the S-wave LVZ was not well constrained. Fig. 3d shows that increasing the minimum LVZ velocity to 4.4 km s -1 has little effect on the synthetic fits to the
49
K. Priestley / Lithos 48 (1999) 45-56
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Fig. 3. (a) Sensitivity test for the depth to the base of the upper mantle lid for the SLR seismogram of the 18 July 1986 earthquake (Fig. 1, No. 2). The solid line is the observed waveform, the dotted line is the synthetic for the southern Africa velocity model of QPM96, and the dashed line is the synthetic for the same velocity model but with the lid base increased to the depth indicated at the left of each seismogram. (b) Same as (a) but for the SLR seismogram of the 10 March 1989 earthquake (Fig. 1, No. 5). (c) Same as (a) but for the SUR seismogram of the 24 July 1991 earthquake (Fig. 1, No. 10). (d) Sensitivity test of the SUR seismogram of the 24 July 1991 earthquake to the minimum S-wave velocity of the LVZ. The solid line is the observed waveform, the dotted line is the synthetic for the southern Africa velocity model, and the dashed line is the synthetic for the same velocity model but with the S-wave velocity of the LVZ increased to the value indicated at the left of each seismogram.
observed waveforms at 2106 km distance, but increasing the minimum velocity to 4.5 km s -1 or greater results in an early higher mode arrival for the synthetic seismogram compared to the observed seismograms at this distance. These tests show that even though the average mantle lid thickness from the waveform inversions is 80 km and the minimum LVZ S-wave velocity is ~ 4.3 km s -1 the thickness of the seismic lithosphere (crust plus upper mantle lid) can be increased to ~ 160 km and the minimum LVZ S-wave velocity can be increased to ~ 4 . 4 5 km s -1 without significantly degrading the match of the synthetic and observed waveforms. Fig. 4 compares synthetic seismograms computed for the revised southern
Africa velocity model with the observed seismograms in three distance ranges. The seismograms at increasing epicentral distance range constrain the increasingly deeper structure in the model. The simultaneous match of the Love (SH) and Rayleigh (P-SV) waveforms with the same isotropic velocity model show that if anisotropy is present, its effects are not sufficiently strong to create a discrepancy between the Love and Rayleigh waves. Increasing the thickness of the seismic lithosphere to 160 km brings the theoretical fundamental mode Rayleigh wave phase velocity curve into agreement with the observed phase velocity curve (Fig. 2) and improves the agreement between the seismically determined velocity structure and the velocity estimates from the
50
K. Priestley / Lithos 48 (1999) 45-56
Fig. 4. Three-component waveform fits at three distance ranges for synthetic seismograms (dotted line) computed from the revised southern African velocity model and the observed seismograms (solid lines). The Love and Rayleigh wave seismograms are fit with the same velocity model, implying that at least the upper mantle lid is isotropic. Event 860718 is Fig. 1, event no. 2, 910724 is event no. 10, and 940818 is event no. 8.
deeper garnet peridotite nodules. The revised southern African shear-wave velocity model is compared in Fig. 5 with the model of QPM96 and the velocities they estimate from the upper mantle nodules. QPM96 examined how uncertainties in earthquake location and source mechanism resulted in uncertainties in their velocity model. The accuracy of the ISC and PDE hypocenters in southern Africa is not known. QPM96 compared the ISC locations of South African rock bursts with locations determined by the South African Geological Survey and estimated the ISC location errors to be about _+ 10 km. They then perturbed the ISC or PDE hypocenter _+ 10 km and reinverted the regional waveforms to evaluate the effects of this magnitude location error on the inversion structure and found that the distance error produced an insignificant change in the inversion model. However, they found that depth errors
resulted in large changes in the synthetic waveforms but that for five of the eight events analyzed, the ISC depth resulted in well-fit waveforms; the other three events required less than a 6 km shift in the focal depth from the ISC or PDE focal depth. The study of Zhao et al. (1997) helps to better evaluate the earthquake location errors. The August 18, 1994 Lake Rukwa earthquake (Event 8, Fig. 1) was well recorded by 20 broadband digital seismographs at distances ranging from 160 to 800 km. Zhao et al. (1997) used teleseismic depth phases to constrain the focal depth to 25 _+ 2 km, similar to the depth (27 • 2 km) Foster and Jackson (1998) found from modeling the teleseismic P- and SH-waveforms. Zhao et al. (1997) then used the Pn arrival times at the regional stations to determine the origin time and epicentral location. They found the origin time for this event to be 00:45:48:79 UTC and the
K. Priestley / Lithos 48 (1999) 45-56 Shear velocity (km/s) 3.25
3.50
3.75
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4.75
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100
x~ 200
300 ~. ---
Revised south Africa Qiu et al (1996) 1 N o d u l e derived estimates
400 Fig. 5. Comparison of the revised shear wave velocity model for southern Africa with the shear wave velocity model of QPM96 and the shear wave velocities estimated in QPM96 from the upper mantle nodules. The thin solid lines denote estimates of the uncertainties in the revised shear wave velocity model from the waveform fitting tests described in the text, the earthquake location errors as described in QPM96, and ~ 2% anisotropy as proposed by Vinnik et al. (1995).
epicenter to be 7.44~ _+4.2 km) S, 31.77~ _+5.3 kin) E, similar to the ISC (00:45:47(_+ 1.4) s, 7.46 ~ (_+0.023 ~ S, 31.71~ +0.029 ~ E, 25 km depth) and PDE (00:45:47.2 UTC, 7.43~ 31.75~ 25 km depth) locations. QPM96 found that the regional waveforms were best fit for a focal depth of 30.9 km. The Zhao et al. (1997) location for the Lake Rukwa earthquake suggests that the +_ 10 km location error assumed by QPM96 was reasonable for at least the later events they analyzed. QPM96 made similar tests for the effects of errors in the CMT source mechanisms and found that _+ 10 ~ variations in the strike, dip and slip made little difference in the inversion velocity structure. QPM96 note that any significant anisotropy in the lithosphere or upper mantle could bias their results. The effect of anisotropy on seismic waves is manifest in two ways" azimuthal anisotropy in which
51
waves of the same type have a wave speed which is a function of azimuth; and polarization anisotropy in which waves of different polarization (eg., Love and Rayleigh w a v e s ) p r o p a g a t e with different wave speeds along the same path. Low-frequency surfacewave observations (Nataf et al., 1984; Montagner, 1996) show weak anisotropy in the upper mantle beneath southern Africa. Vinnik et al. (1995) measured shear-wave splitting at seven sites on the Kaapvaal Craton and found an average delay of 0.9 _+ 0.5 s. From this, they conclude ~ 2% anisotropy localized between 150 and 400 km depth and oriented with the fast direction approximately parallel to the direction of the current plate motion, implying that the material in this depth range has been deformed by plate motion in geologically recent time. A similar result has been observed for Australia (Debayle, 1999). Anderson and Dziewonski (1982) comment that in many cases, isotropic earth models cannot provide a uniform fit to Love and Rayleigh wave data and that in such cases, if the Rayleigh wave data were inverted separately, a pronounced LVZ can occur in the resulting velocity model. They also note that areas such as the Canadian Shield show little or no discrepancy between Love and Rayleigh wave dispersion (Brune and Dorman, 1963). QPM96 used the S n velocities they measured from earthquakes and Sn velocities from refraction (Durrheim and Green, 1992) to constrain the shear wave velocity beneath the Moho. Using this constraint, they simultaneously inverted the Love and Rayleigh regional waveforms for the upper mantle velocity structure. They found that the regional Love and Rayleigh waveforms were fit with a single isotropic velocity structure, suggesting that anisotropy was weak. This is consistent with Bloch et al. (1969) who also found no discrepancy between the Love and Rayleigh waves in southern Africa. Thus, while QPM96 could not make any quantitative statements about the level of polarization anisotropy, there was nothing in their data to indicate the presence of polarization anisotropy and that their results were strongly biased by the isotropic assumption. However, the earthquakes examined in both QPM96 and this study all lie north to northeast of the seismographs and the paths studied sample only a small range of azimuths. As a result, their model may be influenced by azimuthal anisotropy. If
52
K. Priestley / Lithos 48 (1999) 45-56
~ 2% anisotropy exists in the 150-400 km depth range as suggested by Vinnik et al. (1995), this could modify the velocity structure of the LVZ shown in Fig. 5.
3. Comparison with global mantle models for southern Africa Global tomographic models (Suet al., 1995; Masters et al., 1996) show higher than average shear wave velocities extending to depths as great as 500 km beneath the cratons. These features have been cited as evidence for high velocity cratonic roots and as support of the deep tectosphere model (Jordan, 1975). Models for the cratons based on regional seismic data (i.e., Brune and Dorman, 1963; Cara, 1979; QPM96) show a high S-wave velocity upper mantle overlying a S-wave LVZ. In this section the vertical S-wave travel time for the revised southern Africa model is compared to the vertical travel time through the global tomographic models S12WM13 (Su et al., 1995) and S16B30 (Masters et al., 1996) to assess the similarity of the integrated S-wave structure between the base of the crust and 400 km depth in these models. Model S12WM13 (Su et al., 1995) was derived from the inversion of body and mantle waveforms and differential and absolute travel time residuals. The model is expanded to spherical harmonic degree 12 to describe the horizontal variation and to Chebyshev polynomial order 13 to describe the radial variation. The model is given in terms of a relative deviation from PREM (Dziewonski and Anderson, 1981). S 12WM 13 shows higher shear wave velocities than in PREM extending to about 500 km depth beneath southern Africa. An average southern African upper mantle velocity model was determined from S 12WM13 by averaging six velocity profiles computed from S 12WM13 for an area encompassing the whole cratonic region of southern Africa (Fig. 1). Even though S12WM13 shows high shear wave velocities extending over a broad region encompassing southern Africa, only velocity profiles beneath the cratonic region shown in Fig. 1 were averaged. The vertical S-wave travel time between 42 and 400 km depth through the southern African velocity model from S12WM13 is 76.33 s. Making the same
crustal correction as that made for S12WM13 (Woodhouse and Dziewonski, 1984), the vertical S-wave travel time through the revised southern Africa S-wave model is 77.65 s. The mean vertical S-wave travel time for the shields in model S 12WM 13 is 77.26( ___0.82) s. Fig. 6 shows synthetic seismograms computed for a modified S 12WM13 model compared with seismograms of the 14 August 1994 earthquake (Fig. 1, no. 8) recorded at three distance ranges between 2000 and 3000 km across southern Africa. The velocity model consists of the revised seismic lithosphere described in Section 2 to a depth of 160 km and the southern African upper mantle model from S 12WM 13 below 160 km depth. The seismic lithosphere of southern African results in the fit of the fundamental mode surface waves. At these distance ranges, the higher mode surface wave corresponds to the S-wave turning in the mantle transition zone; hence, its arrival time is controlled by the S-wave delay time across the upper mantle. The synthetic mantle S-wave arrives early with respect to the observed mantle S-wave. Lowering the shear wave speed of the upper mantle lid to ~ 4.55 km s-1 as in S12WM13 would improve the match of the synthetic and observed mantle S-wave but would result in a poor fit of the fundamental mode surface S n waves and also disagree with the high frequency travel time observation and the lid velocities determined from the garnet peridotite nodules. Model S16B30 (Masters et al., 1996) was derived from the inversion of absolute and differential travel time residuals, surface wave dispersion and polarization measurements, and mode structure coefficients. The model is parameterized laterally by spherical harmonics expanded up to degree 16 and by 30 natural cubic B-splines in radius. Model S16B30 is specified in terms of the percent deviation from an average spherical velocity model which is not retained in the inversion; hence, the velocity structure cannot be determined and a direct comparison of the vertical S-wave travel times for S16B30 and the southern African model cannot be made. Instead, two indirect comparisons are made. The velocity profile for southern Africa from S16B30 was determined as for S12WM13 except using earth model 1066a (Gilbert and Dziewonski, 1975) as the reference earth model. For this case the vertical S-wave
K. Priestley / Lithos 48 (1999) 45-56
53
SUR
I
m
I
J
i
I
l
I
~
I
~
I
,
I
I
i
,
Fig. 6. Three-componentwaveform fits at three distance ranges for synthetic seismograms(dotted lines) computed from composite velocity model consisting of the seismic lithosphere of the revised southern African model above 160 km depth and the southern Africa upper mantle model derived from the global tomographic model S12WM13 below 160 km depth compared to the observed seismograms(solid lines) of the 14 August 1994 earthquake (Fig. 1, No. 8) recorded at three distance ranges. travel time between 42 and 400 km depth is 77.95 s. Masters et al. (1996) correct their data using the global crustal model of Mooney et al. (1998). Assuming this crustal correction is effectively the same as the crustal correction of Woodhouse and Dziewonski (1984) but using the Mooney et al. (1998) south African crust rather than the crust of PEM-C (Dziewonski et al., 1975), the vertical S-wave travel time for the revised southern Africa upper mantle model is 78.12 s. The mean vertical S-wave travel time for the shields in model S16B30, assuming the average spherical model 1066a is the reference model, is 77.98(+ 0.89) s. The degree of similarity between the models can also be indicated by comparing differential S-wave travel times through models for southern Africa and an old ocean basin. Using the old ocean basin ( > 110 mya) S-wave model of Nishimura and Forsyth (1989), the difference between the vertical S-wave travel time be-
tween 42 and 400 km depth beneath southern Africa and an old ocean basin is - 1.00 s for model S16B30; for the revised southern Africa model it is - 2 . 0 1 s. These comparisons show that the vertical S-wave travel time through the southern Africa shear wave models derived from the regional and global data is similar and lies within the scatter of the shields in each of the global models. However, the models derived from regional seismic data have a high velocity upper mantle lid overlying a shear wave LVZ, whereas the global models have a much smoother velocity structure.
4. Discussion and conclusions The shear wave structure of the QPM96 velocity model for southern Africa has been revised so as to agree with both the fundamental mode Rayleigh
54
K. Priestley / Lithos 48 (1999) 45-56
wave phase velocity data and the regional waveform data. The revised S-wave velocity structure is similar to that of the original QPM96 model except that the high velocity upper mantle lid extends to 160 km depth in the revised model rather than to 120 km depth as in the original model. Increasing the depth to the top of the shear wave LVZ results in fitting of the Rayleigh wave phase velocity data while retaining a satisfactory match to the regional waveform data. The increased lid thickness also improves the agreement between the seismic velocity structure and the velocity estimates at these depths from the garnet peridotite nodules. No revision of the P-wave model of QPM96 is required. The high velocity lid is necessary to fit both the observed high frequency S n arrivals and the propagation of the low frequency fundamental mode Love and Rayleigh waves. The observed velocity over southern Africa is ~ 4.65 km s -1, but higher velocities have been observed in more restricted regions of the Kaapvaal craton (Durrheim and Green, 1992), suggesting that 4.65 km s-1 is a lower bound on the cratonic velocity. These high velocities observed in the upper mantle lid do not result from neglecting anisotropy since the waveform inversion of QPM96 simultaneously fits both the Love and Rayleigh waveforms with a single model. The high S-wave velocities of the lid are further substantiated by their similarity to the S-wave velocities determined from the peridotite nodules from the same depth range. The high lid S-wave velocities require lower velocities beneath the lid to fit the higher mode surface waves. However, the distribution of the low S-wave velocities is not well constrained. The sensitivity tests in Section 2 indicate that the top of the S-wave LVZ is at ~ 160 km depth and the minimum S-wave velocity in the LVZ is 4.30 to 4.45 km s-~ Vinnik et al. (1996) also find low S-wave velocities in the upper mantle beneath southern Africa using receiver function analysis but place the low S-wave velocities in the region starting about 50 km above the 410-km discontinuity. We place the low shear wave velocities at shallow depth beneath the high velocity lid since tests of thicker lid structures result in earlier arriving and larger amplitude lid phases than are observed in the regional seismograms. The vertical S-wave travel times across the top 400 km of the mantle of the revised southern Africa
model are similar to the vertical S-wave travel times through the global tomographic models S12WM13 (Suet al., 1995) and S16B30 (Masters et al., 1996). However, the details of the upper mantle S-wave structure are different in the regional and global models. The shear wave models for southern Africa derived from regionally recorded seismic data (Bloch et al., 1969; Cichowicz and Green, 1992; QPM96) all show high shear wave velocities beneath the Moho with a shear wave LVZ lying below the high velocity lid. The southern African structure from S12WM13 has lower velocities below the Moho, no high velocity lid structure, and no LVZ. All seismological techniques are prone to some averaging when used to determined velocity structure. The velocity model derived from regional seismic data such as that of QPM96 average the velocity structure laterally across the Kaapvaal craton, the Limpopo belt, and the Zimbabwe craton and portions of the surrounding mobile belts. Analysis of the regional seismic data show that the base of the high S-wave velocity upper mantle lid beneath southern Africa lies at ~ 160 km depth. This is an average for the region and the seismic lithosphere is probably somewhat thicker beneath the craton than beneath the mobile belts as suggested by Boyd and Gurney (1986). This depth for the seismic lithosphere is supported by petrological studies of the garnet peridotite nodules and diamond inclusions from the kimberlites (Boyd et al., 1985; Boyd and Gurney, 1986; Boyd, 1989; Gurney, 1990). Modeling of regional waveform data for the Australian shield suggests that the high velocity continental roots there extend to 200-225 km depth (Kennett et al., 1994), similar to that observed in southern Africa but shallower than observed in the global tomographic models for the Australian shield. The similarity of the vertical S-wave travel times in the regional and global models for southern Africa suggest that much of the deep, broad, high velocity root beneath this region in the global tomographic models may be the result of both lateral and vertical averaging. Averaging of the velocity structure is evident in other parts of the global models. For example, both S12WM13 and S16B30 show low S-wave velocity structures associated with the midocean ridges extending to over 300 km depth. It is generally believed that the mid-ocean ridges are
K. Priestley / Lithos 48 (1999) 45-56
passive features caused by lithospheric stretching. Hot asthenospheric material rises to fill the void resulting from the stretching and the magmatism observed along the ridge crest is caused by decompression melting at shallow ( ~ 60 kin) depths (McKenzie and Bickle, 1988). Hence, the low S-wave velocities associated with the mid-ocean ridges are thought to be confined to shallow depths and should not be distinguished in the velocity models to ~ 350 km depth. The high velocity continental roots of the cratons extending to ~ 200 km depth are prone to similar distortion in the global tomographic models. The other difference between the regional and global seismic models is in the way that the crustal effect is handled. In modeling regional seismic data, the effect of the crust is explicitly accounted for in the modeling. In most of the global analysis, a crustal correction is applied to the data before inversion using a standard continent/ocean crustal model or a global crustal model. As Masters et al. (1996) indicate, the crust has a significant effect on most of the global data sets while being too thin to be resolved by them. However, the crustal correction increases the variance of most of the data, which leads to large amplitude velocity perturbations in the uppermost mantle.
Acknowledgements I would like to thank E. Debayle, J. Haines, G. Laske, D. McKenzie, and G. Masters for many helpful discussions and C. Langston, D. McKenzie, H. Patton and R. Saltzer for constructive reviews of the manuscript. I would also like to thank Guy Masters for support while I was on sabbatical leave at the Scripps Institution of Oceanography, where most of this study was undertaken.
References Anderson, D.L., Dziewonski, A.M., 1982. Upper mantle anisotropy: evidence from free oscillations. Geophys. J. R. Astron. Soc. 69, 383-404. Bloch, S., Hales, A.L., 1968. New technique for the determination of surface wave phase velocities. Bull. Seismol. Soc. Am. 58, 1021-1034.
55
Bloch, S., Hales, A.L., Landisman, M., 1969. Velocities in the crust and upper mantle of southern Africa from multi-mode surface wave dispersion. Bull. Seismol. Soc. Am. 59, 15991629. Boyd, F.R., 1989. Compositional distinction between oceanic and cratonic lithosphere. Earth Planet. Sci. Lett. 96, 15-26. Boyd, F.R., Gurney, J.J., 1986. Diamonds and the African lithosphere. Science 232, 472-477. Boyd, F.R., Gurney, J.J., Richardson, S.H., 1985. Evidence for a 150-200-km thick Archaean lithosphere from diamond inclusion thermobarometry. Nature 315, 387-389. Brune, J.N., Dorman, J., 1963. Seismic waves and earth structure in the Canadian shield. Bull. Seismol. Soc. Am. 53, 167-210. Cara, M., 1979. Lateral variation of S velocity in the upper mantle from higher Rayleigh modes. Geophys. J. R. Astron. Soc. 57, 646-670. Cichowicz, A., Green, R., 1992. Tomographic study of uppermantle structure of the South African continent using waveform inversion. Phys. Earth Planet. Inter. 72, 276-285. Debayle, E., 1999. SV-wave azimuthal anisotropy in the Australian upper mantle: preliminary results from automated Rayleigh waveform inversion, Geophys. J. Int., submitted. Durrheim, R.J., Green, R.W.E., 1992. A seismic refraction investigation of the Archaean Kaapvaal Craton, South Africa, using mine tremors as the energy source. Geophys. J. Int. 108, 812-832. Dziewonski, A.M., Anderson, D.L., 1981. Preliminary reference Earth model. Phys. Earth Planet. Inter. 25, 297-356. Dziewonski, A.M., Hales, A.L., Lapwood, E.R., 1975. Parametrically simple Earth models consistent with geophysical data. Phys. Earth Planet. Inter. 10, 12-48. Foster, A.N., Jackson, J.A., 1998. Source parameters of large African earthquakes: implications for crustal rheology and regional kinematics. Geophys. J. Int. 134, 422-448. Fuchs, K., Muller, G., 1971. Computation of synthetic seismograms with the reflectivity method and comparison with observations. Geophys. J. R. Astron. Soc. 23, 417-433. Gilbert, F., Dziewonski, A.M., 1975. An application of normal mode theory to the retrieval of structural parameters and source mechanism from seismic spectra. Philos. Trans. R. Soc. London A 278, 187-269. Gomberg, J.S., Masters, T.G., 1988. Waveform modelling using locked-mode synthetics and differential seismograms: application to determination of structure of Mexico. Geophys. J. Int. 94, 193-218. Gomberg, J.S., Priestley, K., Masters, T.G., Brune, J.N., 1988. The structure of the crust and upper mantle of Northern Mexico. Geophys. J. Int. 94, 1-20. Gurney, J.J., 1990. The diamondiferous roots of our wandering continents. South. Afr. J. Geol. 93, 423-437, 424-437. Jordan, T.H., 1975. The continental tectosphere. Rev. Geophys. 13, 1-12. Kennett, B.L.N., Gudmundsson, O., Tong, C., 1994. The uppermantle S and P velocity structure beneath northern Australia from broad-band observations. Phys. Earth Planet. Inter. 86, 85-98. Masters, T.G., Johnson, S., Laske, G., Bolton, H., 1996. A
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shear-velocity model of the mantle. Philos. Trans. R. Soc. London A 354, 1385-1411. McKenzie, D., Bickle, M.J., 1988. The volume and composition of melt generated by extension of the lithosphere. J. Petrol. 29, 625-679. Montagner, J.-P., 1996. Can seismology tell us anything about convection in the mantle?. Rev. Geophys. 32, 115-137. Mooney, W.D., Laske, G., Masters, T.G., 1998. Crust 5.1: A global crustal model at 5~215 5~ J. Geophys. Res. 103,727-747. Nataf, H.C., Nakanishi, I., Anderson, D.L., 1984. Anisotropy and shear-velocity heterogeneities in the upper mantle. Geophys. Res. Lett. 11, 109-112. Nishimura, C., Forsyth, D., 1989. The anisotropic structure of the upper mantle in the Pacific. Geophys. J. 94, 193-218. Qiu, X., Priestley, K., McKenzie, D., 1996. Average lithospheric structure of southern Africa. Geophys. J. Int. 127, 563-587.
Su, W., Woodward, R.L., Dziewonski, A.M., 1995. Degree 12 model of shear velocity heterogeneity in the mantle. J. Geophys. Res. 99, 6945-6980. Vinnik, L.P., Green, R.W.E., Nicolaysen, L.O., 1995. Recent deformation of the deep continental root beneath southern Africa. Nature 375, 50-52. Vinnik, L.P., Green, R.W.E., Nicolaysen, L.O., Kosarv, G.L., Petersen, N.V., 1996. Deep seismic structure of the Kaapvaal craton. Tectonophysics 262, 67-75. Woodhouse, J.H., Dziewonski, A.M., 1984. Mapping the upper mantle: three-dimensional modeling of Earth structure by inversion of seismic waveforms. J. Geophys. Res. 89, 59535986. Zhao, M., Langston, C.A., Nyblade, A.A., Owens, T.J., 1997. Lower crustal rifting in the Rukwa Graben, East Africa. Geophys. J. Int. 129, 412-420.
LITHOS ELSEVIER
Lithos 48 (1999) 57-80
0
Imaging the continental upper mantle using electromagnetic methods Alan G. Jones Geological Survey of Canada, 615 Booth St., Room 218, Ottawa, Ontario, Canada K1A OE9
Received 27 April 1998; received in revised form 11 November 1998; accepted 20 November 1998
Abstract The internal structure of the continental lithosphere holds the key to its creation and development, and this internal structure can be determined using appropriate seismic and electromagnetic methods. These two are complementary in that the seismic parameters usually represent bulk properties of the rock, whereas electrical conductivity is primarily a function of the connectivity of a minor constituent of the rock matrix, such as the presence of a conducting mineral phase, e.g. carbon in graphite form, or of a fluid phase, e.g. partial melt or volatiles. In particular, conductivity is especially sensitive to the top of the asthenosphere, generally considered to be a region of interconnected partial melt. Knowledge of the geometry of the lithosphere/asthenosphere boundary is important as this boundary partially controls the geodynamic processes that create, modify, and destroy the lithosphere. Accordingly, collocated seismic and electromagnetic experiments result in superior knowledge than would be obtained from using each on its own. This paper describes the state of knowledge of the continental upper mantle obtained primarily from the natural-source magnetotelluric technique, and outlines how hypotheses and models regarding the development of cratonic lithosphere can be tested using deep-probing electromagnetic surveying. The resolution properties of the method show the difficulties that can be encountered if there is conducting material in the crust. Examples of data and interpretations from various regions around the globe are discussed to demonstrate the correlation of electromagnetic and seismic observations of the lithosphere-asthenosphere boundary. Also, the observations from laboratory measurements on candidate mineralogies representative of the mantle, such as olivine, are presented. 9 1999 Elsevier Science B.V. All rights reserved. Keywords: Electrical conductivity; Upper mantle property; Mantle anisotropy
1. Introduction Knowledge of the internal structure of the lithosphere and the geometry of the lithosphere-asthenosphere boundary are critically important for developing our understanding of the dynamics of the Earth. Models of lithospheric growth proposed by different authors are based on differing hypotheses, and many
of these hypotheses are unfettered by constraints. Similarly, geodynamic models of mantle flow generally assume a simplistic geometry for the lithosphere-asthenosphere boundary, whereas topography on this boundary has significant implications for such flow models, as shown in the paper by De Smet et al. (1999, this issue). Deductions about the depthvariation of appropriate physical parameters through-
0024-4937/99/$- see front matter 9 1999 Elsevier Science B.V. All fights reserved. PII: S0024-4937(99)00022-5
58
A.G. Jones/Lithos 48 (1999) 57-80
out the whole of the lithosphere and into the asthenosphere can be used to constrain these hypotheses and models. Evolutionary models of lithospheric growth of continental roots by Jordan (1988) and Ashwal and Burke (1989) appeal to uniformitarianism, and advance modern-day processes as explanations of tectonic events that occurred since the Earth formed. Jordan (1988) suggested that repeated cycles of differentiation and collisional thickening lead to a mantle root, whereas Ashwal and Burke (1989) proposed that cratons are formed by assembling collided island arcs composed of depleted mantle material. In contrast, Helmstaedt and Schulze (1989) suggested more buoyant subduction occurred during the Archean, as a consequence of higher spreading rates, leading to continental roots being formed by imbrication of shallowly subducted slabs. Kusky (1993) modified this model by incorporating trapped wedges of fertile mantle within the stack of imbricated slabs. Models of growth by non-tectonic processes are those of Thompson et al. (1996), who proposed that lithosphere thickened slowly by basal accretion of mobile asthenospheric material, and Polet and Anderson (1995), who hypothesized that permanent roots beneath old cratons may be quite small, and that cold downwelling in the asthenosphere induced by them increases their apparent size and depth. Hoffman (1990) discussed some of these models, and their implications, for the root of the Canadian shield part of the North American craton, concluding that the existing geological data are not adequate for discriminating between the various models, nor can the data define new models that can explain the origin of deep cratonic roots. The apparent Precambrian age of the root is compatible with the Helmstaedt and Schulze (1989) model, but not with either the Jordan (1988) or Ashwal and Burke (1989) result, and is disputed by Thompson et al. (1996) and Polet and Anderson (1995). All of these models have different physical characteristics that can be tested with appropriate geophysical data, and clearly new data are required to further our understanding of the roots of the continents. Unfortunately, there are but two geophysical techniques that can measure physical parameters of the lithosphere. Whereas the density and thermal parameters are inferred using gravity and geothermics, the
seismic wave parameters and electrical conductivity can be measured, albeit through spatially-averaging filters, with appropriate surface-based methods. Since the early 1990s, there has been an explosion in the numbers of teleseismic observations, and some limited attempts using controlled-source seismology, for lithospheric mantle studies. In contrast, the electromagnetic (EM) community was studying the submantle lithosphere in the 1970s and early-1980s, but since the mid-1980s the focus has been more on the crust. This emphasis was due, in the main, to improved instrumentation and time series processing methods for higher frequency data concurrent with global research thrusts and the program directions of the major funding agencies. Recent development of stable and highly-sensitive (30 pT and lower noise levels) magnetometers for long period measurements, coupled with a renewal of scientific interest in the sub-crustal lithosphere, have resulted in a number of high-quality deep-probing EM experiments being carried out in the last few years. The results of most of these though have yet to be published. This paper will briefly review the principles of EM appropriate for imaging the Earth, and describes the natural-source magnetotelluric (MT) method. Some results for the lithosphere and asthenosphere are discussed under the categories of (i) continental lithospheric mantle resistivity, (ii) electrical asthenosphere, (iii) deep mantle observations, and (iv) mantle electrical anisotropy. Laboratory conductivity studies on mantle rocks are presented, together with some discussion of the correlation of electrical and seismic asthenospheres. Conclusions discuss the various hypotheses for lithospheric growth and possible tests using appropriate EM data.
2. Principles of electromagnetism 2.1. Electromagnetic propagation The propagation of an electromagnetic (EM) wave through a uniform homogeneous medium is governed by the propagation constant, k, given by k 2 = o9~( i O"-- WE) where o9 is the radial frequency of oscillation, /z is the magnetic permeability, o- is the electrical con-
A.G. Jones/Lithos 48 (1999) 57-80
ductivity and E is the electric permittivity. Accordingly, the physical parameters being sensed in any electromagnetic experiment are electrical conductivity and electric permittivity (analogous to magnetic susceptibility) within a volume of physical dimension given by an inductive scale length at frequency LO.
For most Earth materials, at frequencies less than a few tens of thousands of Hertz, the contribution from the second term in the equation, w z~E, which describes the displacement currents deduced by James Clerk Maxwell in the 1860s, can be ignored as it is many orders of magnitude smaller than the first term, o~cr, which describes the flow of conduction currents. Once displacement currents are ignored from the wave equations, the equations reduce to familiar diffusion equations met in other branches of geophysics (e.g., gravity, heat flow), with the important difference that they are vector diffusion equations, not scalar ones. General solutions to the vector diffusion equations are given by linear combinations of elementary solutions describing exponential decay within the medium. In a uniform medium, a measure of the inductive scale length is given by the skin depth, S, which is the distance by which the amplitude is 1 / e ' t h of its initial value, given by 6=
[m] ~/~o-
which reduces to S = 0.503v/pT [km] where p is the electrical resistivity (in ~ m), given by 1/o-, and T is the period of oscillation (in s), given by 2rr/w. Accordingly, with electromagnetic methods penetration to all depths is assured from the skin depth phenomenon one merely needs to measure at lower and lower frequency (longer and longer period) to probe deeper and deeper into the Earth. 2.2. Conduction types
There are two dominant types of conduction currents in the crust and upper mantle: electronic conduction and ionic conduction.
59
Electronic conduction (electrons or polarons as charge carriers) is the dominant conduction mechanism in most solid materials, and is a thermallyactivated process governed by the appropriate activation energy for the material, Boltzmann's constant, and the absolute temperature. Ionic conduction (ions as charge carriers) is the dominant conduction mechanism in fluids, but is also important for olivine at high temperatures (11001200~ 2.3. Mental manipulation
The controlling physical parameter in most EM surveys is the electrical conductivity of the medium, o-. Its inverse, electrical resistivity p, is more often discussed in the literature, mainly because the values of conductivity are small fractions, and thus less easy to work with. However, the physical laws are governed by the connectivity of a conducting component within the medium, and so literature from laboratory measurements on rocks is usually in conductivity. Also many of the forward and inverse algorithms are based on the variation of conductivity with depth and lateral distance. This duality is akin to slowness and velocity in seismic studies. An additional complication for the recreational peruser of EM literature is that data are presented in terms of variation with frequency or period, and sometimes both in the same publication, depending on the periodicity of the exciting external wave field and the problem being investigated. Generally, if the depth of interest is in the near surface, then the data will be presented in terms of decreasing frequency, from kiloHertz to Hertz, whereas for deeper studies they will be presented in terms of period, from seconds to thousands of seconds. In this publication period is used exclusively. 2.4. Electrical resistivity
The electrical resistivity of Earth materials varies by more orders of magnitude than any other physical parameter with the exception of viscosity, from one million ohm-metres (106 ~ m) for a competent unfractured batholith to one millionth of an ohmmetre (10 -6 ~1 m) for the most conducting sulphides and graphite. Sea water has a resistivity of 0.3 ~1 m,
60
A.G. Jones/Lithos 48 (1999) 57-80
and highly saline brines can be as low as 0.005 s m (Nesbitt, 1993). Typical ranges of values for certain rock types and for some conducting phases (saline fluid or graphite film) are given in Fig. 1. The bulk electrical resistivity observed from the surface varies over a far smaller range, typically from 1 s m to 104 ~ m, due to the limited resolving power of the diffusive technique. However, even this reduced range of four orders of magnitude makes EM far more sensitive to the presence of anomalies than most other geophysical techniques. The bulk conductivity of a medium is governed by the amount and interconnectivity of the conducting phase, which is usually a very minor constituent of the rock matrix. Archie's Law is often appropriate as a first-order model for the total conductivity of a medium O'm = O'f~7n where o-m and o-f are the conductivities of the bulk medium and the fluid (conducting) phase, respectively, ~7 is the porosity, and the exponent n has a value between 1 and 2. The host rock conductivity, err , is assumed to be sufficiently low to have little effect. Other estimates for the conductivity of a
two-phase mixture are given by serial and parallel network analogues (Madden, 1976, 1983; Bahr, 1997) and by the Hashin and Shtrikman (1963) bounds. These latter bounds find application in many branches of geophysics, including effective transport properties of two-phase mixtures and the pressure dependence of elastic constants, and the Hashin-Shtrikman upper bound is identical to the one of Maxwell (1892) for widely-dispersed spheres embedded in a fluid.
3. The m e t h o d n
Hjelt and Korja (1993) compare and contrast the various electromagnetic methods used for imaging the Earth's crust and upper mantle, stressing the "zooming" ability of EM from lithospheric scale structures (hundreds of kilometres) to small, nearsurface local scale structures (metres length scale). The natural source magnetotelluric (MT) method is the most appropriate electromagnetic technique for probing into the deep lithosphere of the continents. Controlled-source EM surveys require a very large
I
=
T
u
magnetotelluric sounding
,rn
B
~
-I
t
Fig. 1. Resistivity ranges for Earth materials.
I I
I
A.G. Jones / Lithos 48 (1999) 5 7-80
source to penetrate to depths greater than the middle crust, and has only been done on relatively few occasions. Most of these surveys used power lines, such as the 1200-km-long Cabora Bassa powerline in southern Africa before it came into service (Blohm et al., 1977), the Sierra Nevada DC power line (Lienert, 1979), and the 180-km-long line between Finland and Sweden (Kaikkonen et al., 1996). Constable et al. (1984) used a 200-km-long telephone line. A unique deep-probing controlled source was provided by the huge Kibliny magneto-hydrodynamic generator on the Kola peninsula used by the Russians for underwater communication to submarines around the globe (Velikhov et al., 1986, 1987; Heikka et al., 1984). Problems with logistics and the necessity of dealing with source geometry hamper such work for imaging deep into the mantle. In contrast, the MT method uses the time-varying Earth's magnetic field as its source, and depth penetration is assured. The variations are caused by the interaction of the solar plasma with the Earth's magnetosphere. By Faraday's Law of Induction, this varying magnetic field induces an electric current, and the current generates an electric (called "telluric") field in the Earth, and the strength of the
61
telluric field is dependent on the conductivity of the medium. Hence, by observing the magnetic and electric fields simultaneously, and determining their ratios at varying frequencies (equivalent to depths by the skin depth phenomenon), one can derive the conductivity variation with both depth and distance. It can be shown that for a one-dimensional Earth, if one has perfect data at all frequencies there exists only one model that will fit the data (Bailey, 1970; Weidelt, 1972). This uniqueness theorem separates MT from other potential field methods with nonuniqueness as an inherent property. Whereas it is, of course, never possible to obtain perfect data, the existence of this uniqueness theorem propels MT people to make more and more accurate and precise estimates of the Earth responses. On the surface of the Earth, one measures the time variations of the three components of the magnetic field, and the two horizontal components of the Earth's electric field. An example of their variation over a 24-h period is shown in Fig. 2 from a site in northern Canada (within the Yellowknife seismic array) for 12th January, 1997 (UT). Local magnetic midnight (based on geomagnetic coordinates, rather than geographic ones) at Yellowknife is at approxi-
.x]
ExJ Ey I
,.,~............ .~
Fig. 2. Typical magnetotelluric time series. Shown are the time variations of one day of data (12 January 1997, Universal Time) recorded at an MT site at Yellowknife, northern Canada. The scale bar of the three magnetic components represents 300 nT, and the scale bar for the two horizontal electric components represents 400 mV/km.
A.G. Jones / Lithos 48 (1999) 57-80
62
mately 09:00 UT (01:00 local time), when there is a strong negative excursion of Hx. Two hours later is a substorm event rich in long period activity, followed by high signal activity at shorter periods. From the ratios of any two field components in the frequency domain, one can define a complex impedance, Z~y(O~), at radial frequency o~ between, e.g., the northward-directed electric field E~(o) and the perpendicular eastward-directed magnetic field component Hy(w), from Z y(
=
Hy(
'
and the S.I. units of Z are ohms. The ratios of the powers of the fields can be scaled as an apparent resistivity, akin to DC resistivity exploration, 1 Pa ,x y ( 0) ) -- O.)]d5 M y ( o ) )
I '
and the phase lead of the electric field over the magnetic field is given by ~bxy( w ) = tan- 1( E x ( w ) / H y ( t o ) ) . Over a half-space, the apparent resistivity gives the true resistivity of the half space at all frequencies, and the phase is 45 ~ For a multi-layered Earth, the two parameters vary with frequency, and their variation with frequency can be inverted to reveal the structure. Examples of the apparent resistivity and phase curves for some layered Earth models are given below.
4. Continental lithospheric mantle resistivity We know, from ocean-bottom controlled-source EM experiments, that the uppermost oceanic mantle is highly resistive, of the order of 105 ~ m (Cox et al., 1986). This value is consistent with laboratory measurements on dry olivine and peridote. However, estimates of the resistivity of the continental lithospheric mantle are far lower, typically from 80-200 m. A problem with determining this resistivity correctly is that the resistive lithospheric mantle is usually sandwiched between two conducting layers, the lower crust and the electrical asthenosphere, and
thus we can only obtain a minimum bound on its value as in 1D the MT method is generally insensitive to the actual resistivity of resistive layers. Generating the forward response of an Earth model, comprising a reference model with a conducting lower crust and an asthenospheric region (see below for description), and adding a small amount of random scatter and noise (2% to apparent resistivities and 0.58 ~ to phase), we can invert the model data to examine the resolution capabilities in such situations. Fig. 3 shows the data, the theoretical model, and the smoothest model and a seven-layer model that fit the data to a normalized root-meansquare (RMS) misfit of one, i.e., these models neither underfit nor overfit the data and their errors, but fit to within exactly one standard error on average. Clearly, the true resistivity of the uppermost lithospheric mantle is not well resolved, as shown by the gross underestimate of the smooth model just below the base of the crust and by the very large value for the seven-layer model. This lack of resolution is confirmed by singular value decomposition sensitivity analysis (e.g., Jones, 1982) of the seven-layer model which demonstrates that of the 13 parameters (model parameters are the seven layer resistivities and six thicknesses; eigen parameters are combinations of these), nine are well resolved. The weaklyresolved parameters are the resistivity-thickness of the fourth layer (asthenosphere), the resistivity of the first layer (upper crust), and a mixture of model parameters of the third and fifth layers. The worst-resolved parameter is the resistivity of the third layer (the uppermost lithospheric mantle above the asthenosphere), with an error of over an order of magnitude. Thus, there is no resolution in even such high quality data to the actual resistivity of the uppermost mantle we can only set a minimum bound on it. Its thickness, however, is one of the most well-resolved of the model parameters (thirdbest resolved, error of 0.5%), as it separates the two conducting zones and the amount of separation greatly affects the response. This means that although we cannot usually know the actual resistivity of the continental lithospheric mantle, we can know its thickness well, so can estimate the depth to any electrical asthenosphere. Whereas the existence of a conducting lower crust used to be thought of as anomalous, quite to the
A.G. Jones/Lithos 48 (1999) 57-80
63
upper
~
k
3o o
i
asthenosphere
'
If
Fig. 3. Synthetic data (with 2% scatter and error) generated from a theoretical reference model containing a lower crustal conductor and an asthenospheric zone (heavy line). The two other models, one a smooth inversion and the other a 7-layer inversion, fit the synthetic data to an RMS of 1.0.
contrary a resistive lower crust is very rare. Haak and Hutton (1986) define a lower crust with a resistivity in the range of 100-300 l) m as normal. Precambrian regions have a typical lower crustal conductance (conductivity X thickness) of around 20 S, whereas Phanerozoic regions have a conductance more than an order of magnitude higher, 400 S (Jones, 1992). Accordingly, given the almost ubiquitous existence of conducting zones within the crust, defining the actual resistivity of the mantle beneath the continents is a difficult task that still needs attention.
5. Electrical asthenosphere From the earliest MT measurements made on the continents, it was recognised that there had to be a region of high conductivity (low resistivity) at some
depth in the continental upper mantle to explain the long period descending branch of the apparent resistivity curves. Comparison of the depth to this conducting zone with the depth to the seismic low velocity zone associated with the asthenosphere showed generally good agreement as early as 1963 (S~dfim, 1963; Fournier et al., 1963), and studies since have continued to corroborate the correlation (see below). Accordingly, this conducting zone has been termed the electrical asthenosphere. In 1983 an International Union of Geodesy and Geophysics Inter-Association Working Group was established by the International Associations of Seismology and Physics of the Earth's Interior (IASPEI) and Geomagnetism and Aeronomy (IAGA) on Electromagnetic Lithosphere-Asthenosphere Soundings (ELAS) to work with the ELAS group which had been active in IAGA since 1978. Gough (1987) gave an interim report of the activities of the group. A
64
A.G. Jones/Lithos 48 (1999) 57-80
number of investigations in many countries have been carried out under the auspices of ELAS, with one of the most well-known being the EMSLAB (electromagnetic study of the lithosphere, asthenosphere and beyond) study of the Juan de Fuca subduction zone (Booker and Chave, 1989). Based on empirical evidence, Adfim (1976) proposed that the depth to the conducting ELAS layer beneath the continents can be predicted from the formula h = 155q
-
1.46
where h is the depth in km, and q is the heat flow in heat flow units. This formula is inappropriate for some regions (such as the Slave craton in northwestern Canada with a measured heat flow of ~ 1.2 HFU at Yellowknife, but no conducting layer at ~ 120 km), but appears to be reasonably valid for others.
Using ocean-bottom instrumentation, MT measurements in the North Pacific revealed the existence of an electrical asthenospheric layer by the early 1970s (reviewed by Filloux, 1973). Measurements have now been made in most of the world's oceans, and this general result has been confirmed. Data from locations of varying lithospheric age in the North Pacific demonstrate that the top to the asthenospheric conductor deepens as the age of the lithosphere increases (Oldenburg et al., 1984; Tarits, 1986). With a mean conductance value of 5000 S, and a thickness of 100 km taken from seismic information, Vanyan (1984) used these North Pacific data to suggest that the typical resistivity of the oceanic asthenosphere is about 20 f~ m. Fig. 4 shows a comparison of the MT responses expected above a mantle with and without an asthenospheric zone. The two resistivity-depth models are also shown in the figure (on a logarithmic-depth scale). The reference profile has a resistive crust and
E v
f
Fig. 4. The responses of the reference model (heavy lines) compared to those from the same model with additionally an asthenospheric zone.
A.G. Jones/Lithos 48 (1999) 57-80
mantle without any conducting zones. It was obtained by inverting the East European Platform response of Vanyan et al. (1977), assuming 10% scatter and error on the theoretical data, using the Occam smooth inversion of Constable et al. (1987). The model with the asthenospheric layer is the same as the reference profile, but with a 50-km-thick zone of 10 f~ m resistivity (total conductance of 5000 S) from 175-225 km depth extent. The difference in total conductance between the reference model and the asthenospheric model is 4200 S. Clearly the difference between these two models would be apparent in all but the poorest MT data. Sensitivity to the existence of the asthenospheric layer is in the period range 1000-30,000 s for the apparent resistivity data, with the maximum difference at around 5000 s (where the asthenospheric model phase crosses the reference model phase), with 150 f~ m for the reference model compared to 32 f~ m for the asthenospheric model.
65
Commonly, however, there exist conducting zones within the crust which can have the effect of shielding this asthenospheric response. Fig. 5 shows the response for a reference model containing a lower crustal conductor (LCC) of 10 ~ m from 2 0 - 4 0 km depth, for a total of 2000 S conductance, compared to a model including an asthenosphere. The effect of the asthenospheric zone is now more difficult to detect and is at longer periods. At 10,000 s, the difference between the two apparent resistivity data is a factor of three, with 61 f~ m without an asthenosphere compared to 21 f~ m with one, which is over half-an-order of magnitude and should be easily resolvable in high quality data. However, the difficulty resolving the asthenospheric zone occurs because the conductance within the zone can be smeared out over a greater depth range. Synthetic data, with 10% error and scatter in apparent resistivity and equivalent in phase (2.9~ generated from the model containing both a conducting lower crust and as-
01
Fig. 5. The responses of the reference model with a lower crustal conductor (heavy lines) compared to those from the same model with additionally an asthenospheric zone.
66
A.G. Jones/Lithos 48 (1999) 57-80
thenosphere, when inverted using Occam to find the smoothest model fitting the data to a normalized RMS of 1, yields a model without a readily identifiable low resistivity zone representing the asthenosphere (Fig. 6). This screening problem becomes exacerbated if other conducting zones exist within the lithosphere, such as a surficial sedimentary basin or mid-crustal fault zones. As a general rule-of-thumb, we cannot usually sense the existence of an anomalous region if its conductance is less than the total conductance from the surface down to that depth. There appears to be a correlation in the literature between the existence and conductance of electrical asthenospheric zone, and the existence and conductance of a lower crustal conductor. This may be an artifact of the screening effect of the LCC discussed above, and there may not be a tectonic or geody-
namic reason for such a correlation. However, in some situations there may be a causative relationship, such as a thin lithosphere with a correspondingly shallow asthenosphere leading to a partiallymolten, wet lower crust, such as the southern Canadian Cordillera (Jones and Gough, 1995). Selected observations of the electrical asthenosphere are summarized in Table 1, and Praus et al. (1990) gives a comprehensive table of observations in central Europe to 1990. Estimates of the resistivity of the electrical asthenosphere are usually around 5 - 2 0 f~ m. Commonly, this ELAS layer is interpreted as a region of partial melt, given its correlation with the seismically-defined asthenosphere. For a host resistivity of 1000 f~ m, and a melt resistivity of 0.1 ~ m, a melt fraction of 1 - 3 % is sufficient to explain this resistivity for most pore geometries. Partial melt should be gravitationally stable below
/
10
E v
//
1;~000
Fig. 6. Synthetic data (with 10% scatter and error) generatedfrom the reference model with a lower crustal conductor and an asthenospheric zone (heavy lines). The smooth model fits the synthetic data to an RMS of 1.0. The response of the two models are the same to within data error.
A.G. Jones/Lithos 48 (1999) 57-80
67
Table 1 Global observations of the electrical asthenosphere Region
Depth (km)
Resistivity (f~ m)
Reference and Comments
Eastern USSR Kamchatka peninsula
80- > 100
10-20
Sakhalin island
70- > 120 +
5-10
Turanian Shield Siberian Shield
80-120 100
Nizhnyaya Tunguska river Baikal
50- > 100 15- > 25
(Moroz, 1985, 1988) Doming of asth. below centre of peninsula corr. with crustal cond. layer and heat flow maximum (Kosygin et al., 1981) Doming of asth. below island corr. with crustal cond. layer. (Vanyan et al., 1983) No asth. below continent. Sakhalin has low heat flow(1 HFU) (Kharin, 1982 reported in Roberts, 1983) (Safonov et al., 1976; Pavlenkova and Yegorkin, 1983) Corr. with zone of lesser heterogeneity (Safonov et al., 1976) dipping from West-East (Berdichevsky et al., 1980)
20
Eastern Europe Pannonian Basin
50- > 80
Central Europe Rhenish Shield
100- > 150
10
(Bahr, 1985) seismics gives 100 km
Northern Europe Northern Sweden
157-190
2.5-8
Kola peninsula
105
80
West Spitzbergen Karelian megablock
115
1
(Jones, 1980, 1982, 1984; Calcagnile, 1982, 1991; Jones et al., 1983; Calcagnile and Panza, 1987), corr. with sesimic asth. (Vladimirov, 1976) data could not resolve actual resistivity of the layer (Krasnobayeva et al., 1981) Oelsner (1965)
300- > 500
10
(Ritz, 1984) shallower under Phanerozoic than under Precambrian
200 60-150 90 70-150
10 1 20 2-60
(Lilley et al., 1981) band-limited data of very poor quality (Spence and Finlayson, 1983) Whiteley and Pollard (1971) asth. not resovable with these data Moore et al. (1977) asth. barely resolvable with these data
West Africa Senegal Australia Central Australia Eastern Australia
(Adfim et al., 1983) corr. with LVL at 57- > 75 km
200 k m (Agee and Walker, 1993). Lizarralde et al. (1995) offer an alternative hypothesis for the ELAS layer when it is shallower than 200 km, and that is the presence of water, which enhances conductivity by H + ionic conduction (Karato, 1990), dissolved in a depleted, anisotropic, olivine mantle. Free water should not exist in the mantle except at unique tectonic environments, such as in the mantle wedge above a subduction zone (Kohlstedt et al., 1996). Why should the electrical asthenosphere layer be bounded? The M T observations suggest that the electrical asthenosphere layer is of finite thickness extent. Once partial melt starts, why should it cease at greater depths? Perhaps the answer lies in the
shape of the solidus for the mantle mineral and the g e o m e t r y of the geotherm. Figs. 7 and 8 show the solidus for a peridotite that is " w e t " , 0.3 wt.% H 2 0 and 0.7 wt.% CO 2 (Fig. 7a and Fig. 8a) c o m p a r e d to " d r y " , < 0.004 wt.% H 2 0 (Fig. 7b and Fig. 8b) (Olafsson and Eggler, 1983). On Fig. 7a,b are also shown the continental g e o t h e r m of Sclater et al. (1980) and on Fig. 8a,b that of Tozer (1972, 1979). Note that neither g e o t h e r m will explain the existence of an electrical asthenosphere layer for a dry peridotite, but both cross the solidus for a wet peridotite. However, Sclater's geotherm suggests that the electrical asthenosphere layer should be of infinite extent,
68
A.G. Jones/Lithos 48 (1999) 57-80
I
Fig. 7. Peridotite solidus for (a) "wet" (0.3 wt.% H20, 0.7 wt.% CO2) and (b) "dry" (<0.004 wt.% H20) conditions (from Olafsson and Eggler, 1983) compared to the geotherm of Sclater et al. (1980). Note that the geotherm crosses the wet solidus at approx. 35 kbar and does not recross it.
whereas Tozer's geotherm recrosses the solidus to give a layer of limited extent.
uppermost mantle, as the external natural source magnetic fields increase in power with increasing period. However, at periods beyond those that sense the asthenosphere, or around 10,000 s, the power spectrum shows a rapid decrease, with spectral lines at the harmonics of the daily variation (24 h, 12 h, 8 h, 6 h and 4 h are useful). In addition, for a uniform Earth and for the same driving magnetic field power, the electric field decreases in strength with increasing period. This decrease is exacerbated as electrical conductivity decreases with period, i.e., increasing depth, in the Earth. There have been a number of attempts to obtain the Earth's deep conductivity. Most of these using either (i) spherical harmonic analysis of the global observatory observations (e.g., Banks, 1969), (ii) the geomagnetically quiet day source field structure (Sq, e.g., Schmucker, 1985; Olsen, 1 9 9 8 ) o r (iii) long electric dipoles at geomagnetic observatories (e.g., Egbert and Booker, 1992, using Tucson observatory). The response functions from these studies are generally quite poor in that their errors are typically greater than 10%. Also, the one-dimensional (1D) assumption inherent in the application of methods (i) and (ii), and usually adopted in the interpretation of method (iii), is hardly ever shown to be valid.
6. Deep mantle observations Imaging the deep mantle conductivity with reasonable resolution is in its infancy compared to seismological studies. Although there is now compelling seismological evidence for a sharp boundary at the 410 km discontinuity (Vidale et al., 1995), the evidence from EM studies has been equivocal. Indeed, as pointed out by Campbell (1987), models derived from global or large-scale regional EM studies prior to about 1969 usually included an increase in conductivity at 400 km depth, whereas later studies generally did not. Two problems make resolving the mantle resistivity at depths of some hundreds of kilometres difficult. The first is the screening effect of the asthenosphere layer, and the second is due to the natural magnetic field spectrum. Electromagnetic waves are attenuated more strongly in conducting layers than resistive ones, and so one must go to longer and longer periods in order to penetrate through a conducting region. This is not a problem in the crust or
"
Illl
Fig. 8. Peridotite solidus for (a) "wet" (0.3 wt.% H20, 0.7 wt.% CO2) and (b) "dry" (<0.004 wt.% H20) conditions (from Olafsson and Eggler, 1983) compared to the continental geotherm of Tozer (1972, 1979). Note that the geotherm crosses the wet solidus at about 30 kbar and then recrosses it at about 40 kbar. The dry solidus is not crossed by the geotherm.
69
A.G. Jones / Lithos 48 (1999) 57-80
A novel experiment undertaken by Schultz et al. (1993) in a lake in central Ontario, Canada, obtained high quality long period response function estimates by using long electrode lines (1 km) with the electrodes in a chemically- and thermally-stable environment (bottom of the lake). Recordings of the timevarying five-component electromagnetic field were made for over 2 years at the Carty Lake site, and at a second site some 30 km distant the three-component magnetic fields were recorded to use as a "remote reference", familiar in economic theory since the early 1940s (Reiersol, 1950), to remove bias in response function estimation (see, e.g., Gamble et al., 1979). The site was chosen because of its location in the centre of a large Archean craton (Superior province) and from the observed crustal regional electrical homogeneity (Kurtz et al., 1993). The data from the lake site fill the gap between conventional land-based MT responses (to 2 h) and the long period observatory responses (from 6 days). This range was demonstrated to be critical for resolving a peak in conductivity centred at around 450 km (Fig. 9) with a resolving kernel of about 100 km width. Without these lake bottom responses, the peak at 450
km disappears and there is no resolution of the decrease in conductivity below that depth. Using a genetic algorithm approach, Schultz et al. (1993) searched the model space of an overparametrized Earth, and were unable to find acceptable models that excluded an abrupt increase in conductivity near 230, 425 and 660 km, and at these depths conductivity must rise above 0.02 S / m (below 50 1~ m resistivity), 0.06 S / m (16.7 ~ m) and 0.3 S / m (3.3 m), respectively. For western North America, Egbert and Booker (1992) find that a step increase in conductivity at around 400 km is consistent with, but not required by, the estimated responses (Fig. 9; dashed line). A resolvable peak in conductivity at around the 410 km seismic discontinuity is not found in mantle conductivity models from western Europe (e.g., Bahr et al., 1997; Olsen, 1998) although the non-lD aspects of the data suggest that a 1D interpretation may be invalid. Note that the two models for North America, one for the Superior craton and the other for western North America, suggest that the two regions have different physical properties to about 600 km depth. This implies that the craton does indeed have a deep
/
---7 .............
f
I 0"1 i
f
I
.
.
.
.
// / /
/ /
F
0 -'1 "0 cO
o
kP P
Fig. 9. Smooth inversion of the data from the Superior craton (solid line) and from western North America (dashed line). Note the peak in conductivity at around 450 km for the Superior craton, and the rise in conductivity at 400 km for the western North American model (from Schultz et al., 1993 and Egbert and Booker, 1992, respectively).
70
A.G. Jones/Lithos 48 (1999) 57-80
root that may extend beyond the 410 km discontinuity. Recent laboratory experiments present evidence for a step increase in two orders of magnitude in electrical conductivity at the 410 km discontinuity as a consequence of the phase change of olivine to higher pressure mineralogy (wadsleyite) (Xu et al., 1998). A large step increase cannot be resolved using surface data if one obtains a conductivity model with least structure, such as Occam (Constable et al., 1987), even for highly precise data (1% error as shown in Fig. 10b, dotted line). However, with a priori knowledge of the likelihood of a step increase, one can re-perform the inversion without a smooth-
ness penalty at that boundary. The resulting models, for 10% (Fig. 10a) and 1% (Fig. 10b) error data, are consistent with the responses and include the step discontinuity. Note that the models from the responses with higher error are more oscillatory due to a Gibbs-like phenomenon, whereas the step-model for 1% error is almost exact (dashed line, Fig. 10b). With knowledge of this sharp increase in conductivity due to olivine phase change, new high quality data need to be acquired in appropriate regions to test whether the mantle can be effectively described by an olivine mineralogy, or whether the role of volatiles must be invoked. One observation that requires explanation is the observation by Schultz et al. (1993) that mantle conductivity decreases below about 450 km, whereas a phase change from wadsleyite to ringwoodite at 510 km should result in a minor increase in conductivity (Xu et al., 1998).
7. Mantle electrical anisotropy
Fig. 10. Smooth models derived from synthetic data generated from a model with a step transition at 410 km from 0.00333 S / m (300 f~.m) to 0.333 S / m (3 f~ m) (thin line) for (a) 10% error in impedance, and (b) 1% error in impedance. Dotted lines are for models without any discontinuity, whereas the dashed lines are for models with an unpenalised discontinuity at 410 km.
Whereas there has been an explosive growth in the number of observations of seismic SKS shear wave splitting, interpreted as indicative of seismic anisotropy in the lithospheric mantle, there have been very few comparable interpretations of electrical anisotropy. In the main, this has been because the MT community has focused more on using the method to solve crustal-scale problems over the last decade. However, there can also be an ambiguity in the interpretation of MT data in terms of whether the difference between the two modes of propagation is due to 2D s t r u c t u r e , or whether it must be attributed to 2D a n i s o t r o p y on a scale smaller than the resolving power of the technique. For example, the two models in Fig. 11 model A with a lower crustal fault juxtaposing two regions of very different resistivity and model B with a host of 2000 ~ m cross-cut by conducting dykes of 5 f~ m give the same MT response at the observation site (Cull, 1985). The difference between them is in the vertical magnetic field response. Model A will have a large H z response due to the fault, whereas model B will have virtually no measurable H z. So far there has only been one convincing example of a very strong difference in MT response in the
A.G. Jones/Lithos 48 (1999) 57-80
71
Fig. 11. Two models of the lower crust which fit MT observations in Australia. Model A is of a faulted contact in the lower crust, whereas model B is of numerous repetitive resistive dykes (2000 ~ m) cutting a conducting (5 ~ m) lower crust (redrawn from Cull, 1985).
two orthogonal directions without any vertical field response observed, and that is the work of Mareschal et al. (1995) in the Superior Province of Canada. Subsequent SKS work in one region (across the Grenville Front) by Senechal et al. (1996) showed that the fast seismic direction is almost parallel to the high conductivity direction, but there is a small but statistically-meaningful obliquity between them. This obliquity has been interpreted by Ji et al. (1996) as an indicator of the movement sense on ductile mantle shear zones. Support for this hypothesis comes from a recent MT survey across the Great Slave Lake Shear Zone (GSLsz) which yielded anisotropic data at periods sampling upper mantle depths (50 km) with the direction of maximum conductivity at N35E, which is oblique to the strike of the GSLsz of N45E (Jones et al., 1997). The corresponding SKS measurements have yet to be undertaken.
8. Laboratory rocks
conductivity studies on mantle
An informative example of the role of a minor conducting phase is the two-phase mixture H20 ice-KC1 aqueous solution as an analogue for partially molten Earth materials (Watanabe and Kurita, 1993). This system has well-defined phase diagrams and equilibrium structures. When the KC1 concentration is lower than the eutectic concentration (19.7 wt.% at atmospheric pressure), H 2 0 ice and KC1 aqueous solution coexist at a temperature between
the liquidus and the solidus. It also represents a strong contrast in resistivity between solid and melt phases. The variation in conductivity with temperature for 0.2 wt.%, 0.4 wt.% and 0.6 wt.% KC1 is shown in Fig. 12. A dramatic increase in conductivity at the solidus is followed by an increase that can be described by Archie's Law (power exponent of 1.34 to 1.74) to the liquidus. This shows that even at very low melt fractions, much less than 1% melt, the KC1 solution is well interconnected, approaching a connectivity of unity. The conduction process goes from (A) inefficient electronic conduction through the solid ice to (B) efficient ionic conduction in the melt phase at the solidus, to (C) more efficient ionic conduction as the melt fraction rises dramatically at the liquidus (Fig. 13), to (D) completely efficient ionic conduction in the totally liquid phase. Comparisons of the effect of partial melting on physical parameters for this system demonstrate that a melt fraction less than 5% causes electrical conductivity to increase by 1.5-3 orders of magnitude (Fig. 12), whereas the seismic velocity decreases by only a few percent (Watanabe and Kurita, 1994). Laboratory studies on candidate samples of the continental upper mantle have been undertaken for over two decades. Most of the work has concentrated on single crystals of olivine, although some studies have been undertaken on polycrystalline compacts. Reports from different laboratories on various samples were initially very disparate until the role of oxygen fugacity was understood (Duba, 1976), but recent work shows much better agreement (Fig. 14, from Constable et al., 1992). A parametric model
72
A.G. Jones/Lithos 48 (1999) 57-80
1 0 -1 - -
E r~ :~ 10.2 0
o
_
O
•
......
Fig. 12. Conductivity variation with temperature of the two-phase mixture H20 ice-KC1 with differing weight percent KC1 (redrawn from Watanabe and Kurita, 1993). There are four conduction regions. Region A is where there is inefficient electron conduction through the solid ice. Region B and C describe efficient ionic conduction in the melt phase. Region D is completely saturated ionic conduction in the totally liquid phase.
description of the studies is, to a reasonable approximation,
must be other constituents in the rock matrix creating conducting pathways.
O - = 1 0 2 . 4 0 2 e - l . 6 0 e V / k T + 109.17e-4"25eV/kT
in the temperature range 720~176 (Constable et al., 1992). The conduction mechanism is predominantly by electron holes at low temperatures, but then is superseded by magnesium vacancies at high temperatures (Schock et al., 1989). This equation predicts that for a pure olivine mantle, the resistivity at 750~ should be of the order 300,000 f~ m, which is far higher than observed under the one region where we have an excellent estimate of the mantle resistivity due to the lack of any crustal conductors, the Slave craton, and crustal resistivity is in excess of 40,000 f~ m (Jones et al., 1997). The uppermost upper mantle has a resistivity around 4000 f~ m, which would imply a temperature of almost l l00~ which is clearly unreasonable. Accordingly, the upper mantle beneath the Slave craton cannot be pure olivine, and there
Fig. 13. Variation of melt fraction with increasing temperature in the two-phase mixture H20 ice-KC1 (redrawn from Watanabe and Kurita, 1993).
73
A.G. Jones/Lithos 48 (1999) 57-80
l 9
Am | A
9 A A A
9 A A
9 A A
9 A 1
9 A A
1838~ A 9
1727~ 10
Fig. 14. Comparison of olivine laboratory temperature-conductivity data and standard models (Constable and Duba, 1990; Duba et al., 1974; Shankland and Duba, 1990; Tyburczy and Roberts, 1990) (redrawn from Constable et al., 1992).
Laboratory studies of partially molten gabbro by Sato and Ida (1984) and Sato et al. (1986) demonstrate that the volume percent of connected melt increased dramatically in the temperature range of 1127~ to 1200~ from undetectable to 17%. Recent work by Tyburczy and Roberts (1990) shows that melt composition is important for melt conductivity. For the deep mantle, as reported above recent work by Xu et al. (1998) demonstrate that there should be a rapid increase in electrical conductivity at the 410 km boundary as olivine undergoes a phase change to wadsleyite. At 520 km there is a minor increase as wadsleyite transforms to ringwoodite.
9. Correlation of electrical and seismic asthenospheres The existence of a conductive layer within the mantle at the same depth as the low velocity layer
corresponding to the asthenosphere was proposed 35 years ago independently by ,/~dfim (1963) and Fournier et al. (1963). Laboratory studies demonstrate that partial melt will have an effect on both seismic velocities and electrical conductivity (Schmeling, 1985, 1986; Watanabe and Kurita, 1994). An early comparison of asthenospheric zones for the Soviet Union, defined by regions of low velocity and regions of high conductance, shows good spatial correlation (Fig. 15, redrawn from Alekseyev et al., 1977). There has not been a systematic and critical global comparison undertaken to date, but collocated studies demonstrate that this correlation may be common. Calcagnile and Panza (1987) discuss the correlation of results obtained by inversion of surface-wave dispersion data and P-wave travel time observations with those obtained by EM studies for Europe. Praus et al. (1990) compiled data on the MT and seismic determination of the depth to the as-
74
A.G. Jones/Lithos 48 (1999) 57-80
.1"
Q I~~o~~ .
,
..'.
,-,
~-.~..-
Fig. 15. Comparison of regions of high conductance (A) and low velocity zones (B) (redrawn from Alekseyev et al., 1977).
thenosphere in central Europe, and found good correspondence between the two datasets. One cratonic region where there is good correspondence between the base of the lithosphere defined electrically and seismically is the Baltic shield. Early work by Calcagnile (1982), using Rayleighwave dispersion data, gave lithospheric thickness
variation over Fennoscandia shown in Fig. 16a. More recent analysis by Calcagnile (1991), using higher modes, gave a shear wave velocity-depth profile for the path from Copenhagen to Kevo (COP to KEV on Fig. 16a) which exhibits a low velocity zone starting at a depth of around 210 + 30 km and is 20-80 km in thickness extent (Fig. 16b). This model crosses a
75
A.G. Jones/Lithos 48 (1999) 57-80
I,
1 v Q_ 121
l il
~.-
J
300
SAU KIR
~" C?
00
200 I NAT
/
REF
Fig. 16. a) Lithospheric thickness map for Fennoscandia defined by Calcagnile (1982). (b) Shear wave velocity model for the path COP-KEV (Calcagnile, 1991). (c) Resistivity models for regions of the Baltic shield defined by Jones (1980, 1982, 1984) and Jones et al. (1983). range of lithospheric thicknesses in the map of Fig. 16a, and is considered by Calcagnile (1991) to represent the average model for the Baltic shield. Resistivity-depth models derived from EM experiments for four regions of the Baltic shield are shown in Fig. 16c, together with the "reference" model discussed above (REF in Fig. 16c). Models for Kiruna (KIR) and Nattavaara (NAT) were obtained using two dif-
ferent EM methods and using different magnetic source field morphologies. For KIR (and KEV) the horizontal spatial gradient (HSG) approach was used (Jones, 1980, 1982, 1984) on only winter-time highactivity magnetic field time series, whereas for NAT (and SAU) the magnetotelluric method was applied to summer-time low-activity electric and magnetic time series (Jones et al., 1983). Although the HSG
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results have been questioned due to the possible effects of highly non-uniform source fields (Osipova et al., 1989), event-averaging and the analysis approach applied to the HSG data should have mitigated these problems. Also, the correlation of the KIR and NAT models, given the different approaches and disparate datasets, lends credence to their veracity. The published model for NAT was obtained using a Monte-Carlo search approach (Jones and Hutton, 1979b), and gave a lithospheric thickness of 211 + 1 4 / - 1 0 km (Jones et al., 1983). Inverting the NAT data using a modem layered inverse method, the best model has a lithospheric thickness of 213 km with a standard error of + 2 6 / - 23 km. Accordingly, the average depth to the base of the seismologically-defined Baltic shield lithosphere is 210_+ 30 km, and the depth to the base of the electrically-defined lithosphere beneath NAT is 213 _ 24 km. Whilst this remarkable correspondence of the means is far superior than their stated errors, it is clear that there is a statistical correlation between the two. Also, the thinning of the lithosphere to the edges of the Baltic shield, imaged seismically in the map of Fig. 16a, is in agreement with the shallowing of the depth to the electrical asthenosphere beneath KIR and KEV and with the lack of detection of an asthenosphere beneath SAU shallower than about 150 km. (Penetration at SAU is poor because of the lower crustal conducting zone, and also because the response was not defined to sufficiently long periods.) MT measurements at 31 locations in central Finland around SAU show that there is no well-developed asthenosphere, i.e., with conductance greater than 500 S, to 300 km (Korja, 1993).
10. Conclusions Natural-source EM studies of the continents have contributed to our knowledge of their cratonic roots by: 1. Defining the base of the resistive cratonic lithosphere. 2. Imaging the electrical asthenosphere and showing that it correlates with the seismic upper mantle low velocity zone.
3. Demonstrating that there must be some conducting component within the olivine/peridotite matrix that is not present in oceanic mantle lithosphere. 4. Inferring that the continental geotherm must cross the solidus twice to explain the limited depth extent of the electrical asthenosphere layer. Comparisons of the lithospheric thickness derived seismologically and electromagnetically appear to show good correspondence. Thus, there are likely to be significant advances in our understanding when collocated deep seismic and EM experiments are performed, and their datasets jointly inverted. There are many more contributions to be made by deep EM studies. For example, Tozer (1979, 1981) argues that, at the microscopic level, conductivity may be closely related to viscosity. This is highly significant when considering lithospheric processes: conductivity anomalies may be interpreted not only in terms of heat flow anomalies but as zones of possible convection and other transport phenomena. Geometries of the lithosphere-asthenosphere boundary can be imaged to provide data for geodynamic models such as those of De Smet et al. (1998, this issue). Also, hypotheses regarding the lithospheric mantle creation and development, outlined in the Introduction, can be tested using MT data: Hypothesis 1 The cratonic lithosphere is Archean in age and has been strongly coupled to its crust since that time. Test 1 The electrical anisotropy must extend smoothly through the cratonic root and be approximately aligned with surface structures. A more exacting test would require that the obliquity between the seismic and EM anisotropy directions must define the same sense of movement as surface structures (Ji et al., 1996 hypothesis). Hypothesis 2 The cratonic lithosphere is comprised of imbricate stacks formed in the Archean (Helmstaedt and Schulze, 1989 model, with Kusky, 1993 modification). Test 2 There will be a regional variation in electrical anisotropy direction and possibly in electrical parame-
A.G. Jones/Lithos 48 (1999) 57-80
Hypothesis 3
Test 3
Hypothesis 4
Test 4
ters like conductivity or root thickness. The lithosphere has been added to the craton since the end of the Archean by b a s a l accretion ( T h o m p s o n et al., 1996 model). There will be no anisotropy and no regional variation in electrical parameters. If either consistent anisotropy or regional variation are observed, then the hypothesis can be rejected. The old mantle root is of limited spatial extent and its apparent s e i s m i c a l l y - d e t e r m i n e d size is caused by transient cold downwelling (Polet and Anderson, 1995 model). There will be no consistent anisotropy direction, no deep anisotropy, and probably a weak electrical signature over most of the craton. The presence of deep anisotropy and an extensive low conductivity root is cause to reject this hypothesis.
Acknowledgements The author wishes to thank the organizers of the Continental Roots w o r k s h o p for the opportunity to present this material. Discussions with Alan Chave, particularly regarding the hypotheses and tests in the Section 10, and a critical review from him, greatly i m p r o v e d the submitted version. R e v i e w s by S i m o n H a n m e r and Juanjo Ledo, as well as the editorial remarks by Rob van der Hilst, also contributed to the i m p r o v e m e n t s . Geological Survey of C a n a d a Contribution No. 1998257.
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LITHOS ELSEVIER
Lithos 48 (1999) 81-91
Heat flow and the structure of Precambrian lithosphere Andrew A. Nyblade
*
Department of Geosciences, Penn State University, University Park, PA 16802, USA
Received 22 April 1998; received in revised form 7 January 1999; accepted 11 January 1999
Abstract Studies of heat flow from Precambrian terrains have demonstrated three empirical relationships; a temporal relationship between heat flow and tectonic age, a spatial pattern between heat flow and the proximity of Archean cratons, and a temporal relationship between heat flow and the age of lithosphere stabilization. In the first relationship, heat flow is inversely related to tectonic age. The second pattern is characterized by low heat flow from Archean cratons and Proterozoic terrains adjacent to cratonic margins (pericratonic terrains), and higher heat flow from Proterozoic terrains that are more than a few hundred kilometers from a craton. In the third pattern, heat flow decreases as the age of stabilization of the lithosphere increases. A number of interpretations of Precambrian heat flow have been offered to explain one or more of these relationships. The simple cooling of a thermal boundary layer predicts essentially no change in heat flow in terrains older than ~ 1.5 Ga, and therefore does not likely provide a comprehensive framework for the interpretation of Precambrian heat flow. By contrast, two other interpretations, (1) thicker lithosphere beneath Archean terrains than beneath Proterozoic terrains, and (2) greater heat production in Proterozoic crust than in Archean crust, when combined with the special structural configuration of sutures, can both contribute to the spatial and temporal heat flow distributions. Xenolith thermobarometry constraints on lithospheric temperatures, however, limit the contribution of age-dependent crustal heat production, and therefore at least part of the heat flow distributions derive from variations in lithosphere thickness. 9 1999 Elsevier Science B.V. All rights reserved. Keywords: Heat flow; Precambrian; Lithosphere; Xenolith
1. Introduction Heat flow observations from Archean and Proterozoic terrains provide principal constraints on the thermal structure and evolution of Precambrian lithosphere. Studies of heat flow from Precambrian terrains yield three different relationships; (1) a temporal relationship between heat flow and tectonic age
* Tel.: + 1-814-863-8341; fax: + 1-814-863-7823; E-mail: andy @geosc.psu.edu
(Chapman and Furlong, 1977; Sclater et al., 1980), (2) a spatial pattern between heat flow and the proximity of Archean cratons (Ballard et al., 1987; Nyblade et al., 1990) and (3) a temporal relationship between heat flow and the age of lithosphere stabilization (Nyblade and Pollack, 1993a). In this paper, these heat flow relationships are reviewed along with previous interpretations of them. Following this review, the preferred interpretation, which attributes at least part of the heat flow relationships to variations in lithosphere thickness, is then re-examined in light of recent seismic and thermal modeling studies of Precambrian lithosphere.
0024-4937/99/$- see front matter 9 1999 Elsevier Science B.V. All fights reserved. PII: S0024-4937(99)00024-9
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A.A. Nyblade / Lithos 48 (1999) 81-91
2. Precambrian heat flow relationships and interpretations 70
A) The three Precambrian heat flow patterns are illustrated in Fig. 1. The heat flow data used to construct these patterns were taken from the global heat flow data base of Pollack et al. (1993) and comprise 188 observations from eleven Archean terrains (Superior Province in North America; Baltic and Ukrainian Cratons in Europe; Sao Francisco Craton in South America; Kenema-Man Domain, Tanzania Craton and Kalahari Craton in Africa; Karnataka Craton in India; Anabar Craton in Siberia; Pilbara and Yilgran Cratons in Australia) and 675 observations from Proterozoic terrains surrounding these Archean terrains (see Nyblade and Pollack, 1993a for details of the Precambrian heat flow data compilation). These observations come from Precambrian terrains that have not experienced any Phanerozoic tectonism. In constructing Fig. 1, tectonic age was defined as the age of the last major tectonothermal event to have affected the crust at a heat flow site, and is based on dates from isotopic systems (e.g., K-Ar, Ar-Ar and Rb-Sr) that are reset during thermal events which raise rock temperatures to a few hundred degrees. Dates from isotopic systems with higher closure temperatures (e.g., U-Pb and Sm-Nd) were not used to determine tectonic age because such ages may "see through" younger tectonothermal events and reveal older, higher temperature events, including the extraction of crust from the mantle. Tectonic age assignments were made using the Geological World Atlas (Unesco, 1976) in conjunction with regional geological maps, and the data were divided into three age groups; Archean ( > 2500 Ma), Early Proterozoic (2500-1700 Ma) and Late Proterozoic (1700-600 Ma).
2.1. Heat flow-tectonic age relationship The temporal pattern between heat flow and tectonic age is characterized by a decline in the mean heat flow of the age groups with increasing tectonic age (Fig. l a). Statistical " t " tests indicate that the Archean mean heat flow is marginally but significantly different from the Early and Late Proterozoic mean heat flows at a 95% confidence level, and that
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Fig. 1. Precambrian heat flow patterns (after Nyblade and Pollack, 1993a). (a) Heat flow vs. tectonic age. The solid circles show the mean heat flow for each grouping, and the thick and thin vertical error bars give the standard error and standard deviation of the mean, respectively. The horizontal error bars represent the age range for each age group. (b) Heat flow vs. distance from mapped cratonic margins. Measurements to the left of the cratonic margin are from Archean ( > 2500 Ma) terrains, and those to the right are from Proterozoic (600-2500 Ma) terrains. The solid circles show the mean heat flow for 200 km wide intervals and are plotted in the middle of each 200 km interval. The thick and thin vertical error bars give the standard error and standard deviation of the mean, respectively. (c) Heat flow vs. lithosphere stabilization age. The solid circles show the mean heat flow for each grouping, and the thick and thin vertical error bars give the standard error and standard deviation of the mean heat flow, respectively. LPr = Late Proterozoic; EPr -- Early Proterozoic; Ar = Archean.
A.A. Nyblade/ Lithos 48 (1999) 81-91
the Early Proterozoic mean heat flow is statistically different from the Late Proterozoic mean heat flow at a 90% confidence level. In a study by Rao et al. (1982), the validity of a heat flow-tectonic age relationship was questioned because statistical tests suggested that at a 95% confidence level the data within the groupings were not normally distributed. However, in comparing the means of different data groupings, it is not imperative that the data distributions be normal, as along as the distributions show similar variance (e.g., Davis, 1986). Therefore, contrary to the Rao et al. study, it can be argued that Precambrian heat flow shows a valid relationship between heat flow and tectonic age in a general way, as was suggested in earlier studies by Chapman and Furlong (1977) and Sclater et al. (1980).
2.2. Spatial heat flow relationship The spatial heat flow pattern is shown in Fig. lb. Generally, heat flow from Archean cratons and Proterozoic terrains within a few hundred kilometers of an Archean craton margin is lower than heat flow from Proterozoic terrains which lie beyond a few hundred kilometers of a cratonic margin. The mean heat flow in the Archean cratons and Proterozoic terrains within 400 km of a craton margin is between 40-44 mW m -2, and increases to between 54 and 58 mW m -2 in Proterozoic terrains at distances greater than about 400 km from a cratonic margin.
2.3. Heat flow-lithosphere age relationship The temporal relationship between heat flow and the age of lithosphere stabilization is shown in Fig. l c and comes from an examination by Nyblade and Pollack (1993a) of whether or not previous interpretations of Precambrian heat flow could reasonably account for both the temporal and spatial heat flow relationships shown in Fig. l a and b.
2.3.1. Heat flow and a cooling-half space The first interpretation of Precambrian heat flow considered by Nyblade and Pollack was the simple cooling of a thermal boundary layer (Crough and Thompson, 1976). This interpretation was proposed to explain the heat flow-tectonic age relationship, but it cannot account for either the temporal or
83
spatial Precambrian heat flow patterns because it predicts essentially no change in heat flow for terrains older than ~ 1.5 Ga (mid-Proterozoic). The cooling half-space model is not in agreement with the heat flow observations and was rejected by Nyblade and Pollack as providing a comprehensive interpretation of Precambrian heat flow.
2.3.2. Heat flow and lithospheric thickness The second interpretation for Precambrian heat flow examined by Nyblade and Pollack derived from the model proposed by Ballard and Pollack (1987) for the thermal structure of the lithosphere in southern Africa in which heat flow is inversely related to lithosphere thickness. Model calculations of heat flow through a conductive lid (lithosphere) of variable thickness show greater heat flow through thinner regions and lesser heat flow through thicker regions, a pattern arising from the greater thermal resistance of the thick lithosphere (Ballard and Pollack, 1987; Nyblade and Pollack, 1993b). With variations in lithosphere thickness, Nyblade and Pollack argued that it is possible to account for both the spatial and temporal Precambrian heat flow patterns, but it is necessary to invoke a characteristic structural relationship for the sutures at the margins of Archean cratons. Fig. 2 illustrates this structural interpretation and its relationship to the spatial pattern between heat flow and the proximity of Archean cratons; thicker Archean mantle lithosphere lies beneath the region of low heat flow from the Archean craton and adjacent Early a n d / o r Late Proterozoic terrains, and thinner Proterozoic mantle lithosphere lies beneath the region of higher heat flow in the Early and Late Proterozoic terrains further from the mapped craton margin. Corroborating evidence for the special structural configuration of sutures in Fig. 2 comes from geological and geophysical field studies and was summarized by Nyblade and Pollack (1993a). The variability in the extent of thrusting within a n d / o r between Archean and Proterozoic sutures can be estimated from the extent of thrusting observed in Phanerozoic orogenic belts. For example, in the Mesozoic thrust belt of the North American Cordillera the extent of thrusting varies from 100 km in northern Idaho and southern Alberta to over 300 km in parts of the thrust belt further north in Canada
A.A. Nyblade /Lithos 48 (1999) 81-91
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(Price, 1981; Smith, 1981; Wheeler, 1987). Similarly, in the Appalachian orogen the extent of thrusting varies along strike from about 100 km in the New York Promontory to over 400 km in the Pennsylvania and Tennessee Reentrants (Hatcher, 1981; Hatcher et al., 1990). These examples suggest that variations of several hundred kilometers are likely in the extent of thrusting within and between different orogenic belts. This variability in the extent of overthrusting means that the model in Fig. 2, which shows Proterozoic crust overthrust 400 km onto Archean lithosphere, should be thought of as schematic and general. The heat flow-tectonic age relationship (Fig. l a) can also be explained by the lithospheric structure shown in Fig. 2, when one considers the geographic distribution of heat flow sites in Early vs. Late Proterozoic terrains. Fig. 3 shows that with respect to
the total number of measurements in the Early and Late Proterozoic age groups, proportionally more heat flow measurements from Early Proterozoic terrains lie closer to a craton margin, where the heat flow is lower and the lithosphere is thicker, than do measurements from Late Proterozoic terrains. Of the measurements in Early Proterozoic terrains 63% are from sites within 400 km of a cratonic margin, whereas only 46% of the measurements from Late Proterozoic terrains lie within 400 km of an Archean craton. Hence, the mean heat flow for Early Proterozoic terrains appears to be slightly lower than the mean heat flow for Late Proterozoic terrains, yet somewhat higher than the mean heat flow for Archean terrains. The greater percentage of measurements in Early Proterozoic terrains nearer to a cratonic margin is not unexpected, since Early Proterozoic terrains generally lie closer to Archean cratons than do Late
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Proterozoic terrains. Because both the heat flowtectonic age relationship and the spatial pattern between heat flow and the proximity of Archean cratons can be explained by the same model (Fig. 2), Nyblade and Pollack (1993a) suggested that a model employing variations in thickness between Archean and Proterozoic lithosphere coupled with the recognition of the allochthonous nature of suturing between Archean and Proterozoic terrains provided a viable comprehensive interpretation of Precambrian heat flow.
2.3.3. Heat flow and crustal heat production The third interpretation of Precambrian heat flow examined by Nyblade and Pollack is that heat production is greater in Proterozoic crust than in Archean crust (England and Richardson, 1980; Vitorello and Pollack, 1980; Sclater et al., 1980; Morgan, 1984, 1985). A number of models have been proposed to explain why crustal heat production may be greater in Proterozoic crust, such as (1) lesser erosion, (2) greater enrichment of U, Th and K isotopes in the upper crust due to the upward mobilization of incompatible elements during successive orogenic events, (3) secular changes in processes which generate continental crust, and (4) selective survival of low heat production Archean crust (e.g., England and Richardson, 1980; Morgan, 1984, 1985; Taylor and McLennan, 1985). Morgan (1984, 1985) and Morgan and Sass (1984) summarize geochemical
85
evidence suggesting that the mean heat production of Proterozoic crust may be greater than that of Archean crust, and they also argue that models (3) and (4) above are more tenable geologically. Combined with the characteristic structural configuration of sutures shown in Fig. 2, age-dependent crustal heat production may also, at least in principle, be able to account for the Precambrian heat flow patterns. In such an interpretation, the low cratonic heat flow results from low Archean crustal heat production, the higher heat flow in the Proterozoic terrains away from the cratonic margins results from higher Proterozoic crustal heat production, and the low heat flow in the pericratonic terrains results from Archean crust underlying the pericratonic terrains. Such an interpretation requires that Proterozoic crust comprising the overlying thrust sheets in the suture adds little to the surface heat flow. This situation could arise if the thrust sheets are very thin or if the Proterozoic crustal material in the thrust sheets consists of low heat production middle and/or lower crustal rocks (i.e., higher heat production upper crust within the thrust sheets has been eroded away). With differences in Archean and Proterozoic crustal heat production, the temporal Precambrian heat flow pattern can be accounted for by the geographic distribution of heat flow measurements in Early vs. Late Proterozoic terrains described previously. Thus, differences in Archean and Proterozoic crustal heat production might also provide a viable interpretation of Precambrian heat flow.
2.3.4. Heat flow and the age of lithosphere stabilization To summarize, Nyblade and Pollack (1993a) reasoned that both variations in Archean and Proterozoic lithosphere thickness and differences in Archean and Proterozoic crustal heat production, when combined with the special suture geometry illustrated in Fig. 2, could possibly explain the temporal and spatial heat flow patterns. However, Nyblade and Pollack argued further that a significant part of the Precambrian heat flow distributions likely derives from differences between Archean and Proterozoic lithosphere thickness, based on the thermal modeling results of Ballard and Pollack (1987) and Nyblade and Pollack (1993b). These thermal models attempted to satisfy the spatial heat flow pattern in
86
A.A. Nyblade/ Lithos 48 (1999) 81-91
addition to xenolith thermobarometry constraints on mantle temperatures (Boyd et al., 1985; Boyd and Gurney, 1986). From these modeling studies, it was concluded that the heat flow and thermobarometry constraints could be met with variations in lithospheric thickness alone, or by a combination of variations in lithospheric thickness and differences in crustal heat production, but not by differences in crustal heat production alone. Interpretations explaining the heat flow observations with variations only in crustal heat production, while able to generate suitable contrasts in surface heat flow, failed to satisfy temperature constraints in the mantle imposed by xenoliths. In continuing their line of reasoning, Nyblade and Pollack argued that if Precambrian terrains with low heat flow are underlain by thicker lithosphere that stabilized during the Archean, and terrains with higher heat flow are underlain by thinner lithosphere that stabilized during the Proterozoic, as depicted in Fig. 2, then a different heat flow-age relationship for the Precambrian could be proposed, an inverse relationship between heat flow and the age of stabilization of the lithosphere (Fig. l c). Fig. l c was constructed by removing from both the Early and Late Proterozoic tectonic age data groups those points within 400 km of an Archean craton margin, (i.e., those that overlay thick lithosphere stabilized in the Archean), and reassigning them to the Archean group. Because the heat flow at those points are typically in the lower part of the Proterozoic heat flow ranges, but similar to points within the craton, their reassignment raises the means of the Early and Late Proterozoic groups from which they were removed to 57 and 55 mW m -z, but leaves the Archean mean heat flow essentially unchanged at 41 mW m -2. The small difference in the mean heat flow between Early and Late Proterozoic stabilization age groups probably does not have any statistical significance (Nyblade and Pollack, 1993a). In the heat flow-tectonic age relationship (Fig. l a), the margin of the Archean craton is taken as the edge of the mapped Archean crust, the tectonothermal craton margin (see Fig. 2). In the relationship between heat flow and the age of stabilization of the lithosphere (Fig. l c), the craton margin is defined as the edge of the Archean mantle lithosphere, the lithospheric craton margin (see Fig. 2).
A scenario for the stabilization of continental lithosphere suggests why lithosphere thickness and, consequently, heat flow might vary as a function of the age of lithosphere stabilization. The stabilization of cratonic lithosphere has been postulated to result from two processes associated with the extraction of continental crust from the mantle: (1) depletion of basaltic components (e.g., A120 3, FeO, CaO) from the fertile mantle, imparting a chemical buoyancy to the sub-crustal lithosphere (Jordan, 1978, 1979, 1981, 1988), and (2) removal of volatiles from the mantle, which stiffens the residuum and renders it less susceptible to entrainment in mantle convection (Pollack, 1986). Pollack (1986) further argues that the efficiency of the devolatilization and chemical depletion processes is linked to the fractional extent and depth range of partial melting in the mantle, and that both of these quantities were greatest in the Archean when mantle temperatures were higher. Hence, the processes which stabilize continental lithosphere have become less efficient with time, and Proterozoic lithosphere should therefore be thinner than Archean lithosphere, thus providing a simple framework for understanding why heat flow from Precambrian terrains might vary with the age of stabilization of the lithosphere.
3. Discussion
In summary, the preferred interpretation of Nyblade and Pollack (1993a) is a model in which either some or all of the heat flow differences between Archean and Proterozoic terrains arises from the special structural configuration of sutures shown schematically in Fig. 2, and thicker Archean than Proterozoic lithosphere. In this section, this interpretation is re-examined in light of some old and new seismic studies on lithospheric structure, as well as some recent thermal modeling studies. One way to evaluate the Nyblade and Pollack model is to determine if the thickness of Precambrian lithosphere imaged seismically correlates with the heat flow pattern shown in Fig. 2. This evaluation is not necessarily straightforward to carry out, since ascertaining the thickness of continental lithosphere seismically is not simple to do. Low velocity zones within the upper mantle beneath continents
A.A. Nyblade/ Lithos 48 (1999) 81-91
may not be well developed or easily identified. Nevertheless, seismic tomography images of deep continental structure are improving as new imaging techniques are developed and larger data sets become available. In addition, long-range seismic refraction profiles crossing a number of Precambrian terrains in several regions have been recently published. Figs. 4 and 5 show heat flow and seismic models from two areas where there is sufficient data to compare regional heat flow patterns with seismic estimates of the thickness of the Precambrian lithosphere. The first area is the Baltic shield, and the seismic constraints on lithospheric thickness come from the long-range FENNOLORA seismic refraction profile (Guggisburg and Berthelsen, 1987; Fig. 4). All parts of the model in Fig. 2 are well represented by heat flow and the seismic structure of the lithosphere beneath the Baltic Shield. The seismic profile shows thick ( ~ 220 km) lithosphere beneath
9
87
the low heat flow cratonic and pericratonic terrains, and 'thinner lithosphere ( ~ 110 km) beneath the higher heat flow Proterozoic terrains away from the cratonic margin, thus corroborating the Nyblade and Pollack model. The second area with sufficient data to compare heat flow and seismic models is North America. The seismic constraints on lithospheric thickness used here come from a recent shear-wave tomography study by van der Lee and Nolet (1997), where fundamental and higher-mode Rayleigh waveforms from broadband seismograms were inverted to obtain an image of the upper mantle S velocity structure to a depth of 660 km. Among the well-resolved features in this model is a high-velocity keel beneath the North American craton which extends to a depth of about 250 km. Shown in Fig. 5 is a horizontal slice through this model at 170 km depth, illustrating the fast velocities of the cratonic keel, along with heat flow from Precambrian terrains in North America
45
35
Fig. 4. (a) N E - S W heat flow profile across the Baltic Shield. The data come from the global compilation of Pollack et al. (1993). The solid circles show the mean heat flow for each grouping, the vertical error bars give the standard error of the mean heat flow, and the horizontal error bars show the range for each grouping. (b) Seismically-defined lithospheric thickness beneath the Baltic Shield along the same profile (Guggisburg and Berthelsen, 1987).
A.A. Nyblade /Lithos 48 (1999) 81-91
88
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Fig. 5. Heat flow from Precambrian terrains in North America plotted with S wave velocity variations at a depth of 170 km. The S wave model (Na95) is from van der Lee and Nolet (1997). The velocity scale is non-linear to highlight the fast velocities associated with the cratonic keel.
not disturbed by any Phanerozoic tectonism. The heat flow data shown are from an updated version of the Nyblade and Pollack (1993a) compilation. Fig. 5 illustrates that most of the high heat flow observations lie along the periphery of the seismically defined cratonic keel, as predicted by the Nyblade and Pollack model (i.e., Fig. 2). It is also apparent from Fig. 5 that the seismically defined cratonic keel is probably not present beneath the Grenville Province where heat flow is low, and this observation has been used by Jaupart and Mareschal (1999) (this volume) to argue against the existence of Precambrian heat flow patterns. However, this observation does not necessarily invalidate
the heat flow patterns described earlier nor the Nyblade and Pollack interpretation of these patterns, since the absence of thick lithosphere beneath the low heat flow Grenville Province can be explained by destruction of the lithosphere during the last 200 Ma. van der Lee and Nolet (1997) comment on the absence of fast upper mantle velocities beneath the Grenville Province and attribute the "missing" lithosphere to a two-stage process. First, they propose that the lithospheric keel was rehydrated during subduction of the Iapetus plate. Then at 180 and 100 Ma, the Grenville Province passed over the Great Meteror and Monteregian hotspots, respectively (Morgan, 1983). The increased upper mantle temper-
A.A. Nyblade / Lithos 48 (1999) 81-91
atures associated with the plumes melted some of the rehydrated cratonic lithosphere beneath the Grenville Province. If this interpretation of why the lithospheric keel is missing beneath the Grenville Province is correct, then, given the slow rate at which heat conducts through the lithosphere, it can explain why heat flow from the Grenville Province is still low. Fig. 6 shows the transient response of surface heat flow for a sudden change in temperature at the base of a slab, as would be experienced by the base of the Grenville Province lithosphere while passing over a hotspot. If At = 300 K represents a reasonable upper limit to the size of the temperature perturbation (McKenzie et al., 1974; Gurnis, 1988; Schilling, 1991; Gurnis and Zhong, 1991), then Fig. 6 shows that for 100-150 km thick lithosphere (i.e., the thickness of the Grenville Province lithosphere after disruption), surface heat flow will increase by only 5-8 mW m -2 over 180 Ma, an increase that is not easily detectable. Hence, the thermal perturbation to the Grenville Province lithosphere associated with its
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100
200
300
400
Fig. 6. The change in surface heat flow plotted against lithosphere thickness for a 300 K step change in temperature at the base of the lithosphere. The numbers on the curves give the time in Ma elapsed since the step change in temperature at the base of the lithosphere occurred. A thermal conductivity of 3 Wm-1 K-1 and a thermal diffusivity of 1 • 10 -6 m 2 s-z were used in this calculation.
89
disruption may not have yet affected surface heat flow significantly. In addition to the work of Jaupart and Mareschal (1999) (this volume), Lenardic (1997) also questioned the conclusion reached by Nyblade and Pollack (1993a) that at least some part of the Precambrian heat flow patterns arise from differences in thickness between Archean and post-Archean lithosphere. An examination of Lenardic's study helps to clarify some important aspects of the Nyblade and Pollack model (Fig. 2). In his study, Lenardic (1997), using numerical thermal models that allow continents (chemical boundary layers) to form over a convecting mantle, argued that the thickness of continental lithosphere does not significantly influence surface heat flow. Lenardic's argument is based on modeling results suggesting that if a continent with variable lithosphere thickness sits over a downwelling in the mantle, then the surface heat flow across the continent exhibits little variability (see Fig. 8 in Lenardic, 1997). However, Lenardic's models also show that if the same continent sits over an upwelling, then surface heat flow across the continent does vary in accord with the Nyblade and Pollack model. Lenardic's model is different from the model of Ballard and Pollack (1987) in that Ballard and Pollack used a parameterized thermal convection algorithm rather than an algorithm for full thermal convection. In the Nyblade and Pollack (1993b) model, an algorithm for full thermal convection was used but the lithosphere was imposed a priori in the model. Because of these differences in the way the models were constructed, Lenardic pointed out that in the Ballard and Pollack (1987) and Nyblade and Pollack (1993b) models the thinner mobile belt lithosphere sat over mantle upwellings, thus enhancing the delivery of deep mantle heat to the mobile belt lithosphere with respect to the thicker cratonic lithosphere. Because of the dynamic feedback between plates and mantle convection, continents probably do not sit for protracted geological periods over either mantle upwellings or downwellings. Models with dynamically interacting plates and convection show that plates translate laterally as heat builds up beneath them, and this lateral plate motion modulates the time-averaged mantle temperature at the base of the lithosphere by moving the plate from warmer to
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colder regions, where the plates sit until mantle heat builds back up beneath them (e.g., Gurnis, 1988; Gurnis and Zhong, 1991). The models of Ballard and Pollack (1987) and Nyblade and Pollack (1993b) therefore likely represent end-member models which provide an estimate of the maximum difference between heat flow from cratonic and mobile belt lithosphere, while the model of Lenardic (1997) likely represents the other end-member model and provides an estimate of the minimum difference. Since continents probably do not sit over upwellings or downwellings for long periods of geologic time, the amount of heat flow "diverted" by the cratonic keel is in-between that predicted by Ballard and Pollack (1987) and Lenardic (1997). Finally, the conclusion that at least some of the lower heat flow in Archean regions arises from thicker lithosphere is supported by recent analyses of pressure-temperature ( P - T ) estimates from mantle xenoliths found in Archean cratons (Rudnick et al., 1998; Rudnick and Nyblade, 1999). In these analyses, P - T data from the Kalahari Craton were combined with estimates of surface heat flow and modeled using 1D geotherm calculations to place constraints on average crustal heat production, average mantle lithosphere heat production, and lithosphere thickness. The analyses yielded estimates of 0.5-0.8 ~W m -3, 0-0.07 ~W m -3 and 200-250 km for mean crustal heat production, mean mantle lithosphere heat production and lithosphere thickness, respectively. These estimates are consistent with the modeling results of Ballard and Pollack (1987), which show that if the cratonic lithosphere is 200-250 km thick and mean Archean crustal heat production is ~ 0.7-0.8 ~W m -g, then at least some of the difference in heat flow between the Archean and postArchean terrains arises from the presence of the thick cratonic lithosphere. In this case, some of the heat flow differences also arise from variations in mean crustal heat production between the Archean and post-Archean crusts.
4. Summary Studies of heat flow from Precambrian terrains have demonstrated three empirical relationships; a temporal relationship between heat flow and tectonic
age, a spatial pattern between heat flow and the proximity of Archean cratons, and a temporal relationship between heat flow and the age of lithosphere stabilization. A number of interpretations of Precambrian heat flow have been offered to explain one or more of these relationships. The simple cooling of a thermal boundary layer predicts essentially no change in heat flow in terrains older than ~ 1.5 Ga and, therefore, does not likely provide a comprehensive framework for the interpretation of Precambrian heat flow. Two other interpretations, (1) thicker lithosphere beneath Archean terrains than beneath Proterozoic terrains, and (2) greater heat production in Proterozoic crust than in Archean crust, when combined with the special structural configuration of sutures, can both contribute to the spatial and temporal heat flow distributions. Xenolith thermobarometry constraints on lithospheric temperatures, however, limit the contribution of age-dependent crustal heat production, and thus at least part of the heat flow distributions likely derive from variations in lithosphere thickness. This position is supported by seismic images of cratonic lithosphere.
Acknowledgements The author is grateful to S. van der Lee for providing the S wave model of North America shown in Fig. 5 and to Bill McDonough for his constructive review. This research was supported in part by the National Science Foundation grant EAR-9103065.
References Ballard, S., Pollack, H.N., Skinner, N.J., 1987. Terrestrial heat flow in Botswana and Namibia. J. Geophys. Res. 92, 62916300. Ballard, S., Pollack, H.N., 1987. Diversion of heat by Archean cratons: a model for southern Africa. Earth Planet. Sci. Lett. 85, 253-264. Boyd, F.R., Gurney, J.J., 1986. Diamonds and the African lithosphere. Science 232, 472-477. Boyd, F.R., Gurney, J.J., Richardson, S.H., 1985. Evidence for a 150-200 km thick Archean lithosphere from diamond inclusion thermobarometry. Nature 315, 387-389. Chapman, D.S., Furlong, K.P., 1977. Continental heat flow-age relationships (abstract). EOS, Trans. AGU 58, 1240.
A.A. Nyblade / Lithos 48 (1999) 81-91 Crough, S.T., Thompson, G.A., 1976. Thermal model of continental lithosphere. J. Geophys. Res. 81, 4857-4862. England, P.C., Richardson, S., 1980. Erosion and the age-dependence of continental heat flow. Geophys. J. R. Astron. Soc. 62, 421-438. Guggisburg, B., Berthelsen, A., 1987. A two-dimensional velocity model for the lithosphere beneath the Baltic Shield and its possible tectonic significance. Terra Cognita 7, 631-638. Gurnis, M., 1988. Large-scale mantle convection and the aggregation and dispersal of supercontinents. Nature 332, 695-699. Gurnis, M., Zhong, S., 1991. Generation of long wavelength heterogeneity in the mantle by the dynamic interaction between plates and convection. Geophys. Res. Lett. 18, 581-584. Hatcher, R.D., 1981. Thrusts and nappes in the North American Appalachian Orogen. In: McClay, K.R., Price, N.J. (Eds.), Thrust and Nappe Tectonics. Geol. Soc. London Spec. Publ., Geol. Soc. London, London, Vol. 9, pp. 427-488. Hatcher, R.D., Osberg, P.H., Drake, A.A., Robinson, P., Thomas, W.A., 1990. Tectonic map of the US Appalachians, IGCP Project 27. Geol. Soc. Am. Jaupart, C., Mareschal, J.C., 1999. The thermal structure and thickness of continental roots. In: McDonough, W.F., van der Hilst, R. (Eds.), Lithos. Special volume on Continental Roots. Jordan, T.H., 1978. Composition and development of the continental tectosphere. Nature 274, 544-548. Jordan, T.H., 1979. Mineralogies, densities, and seismic velocities of garnet lherzolites and their geophysical implications. In: Boyd, F.R., Meyer, H.O.A. (Eds.), The Mantle Sample: Inclusions in Kimberlites and Other Volcanics, AGU, Washington, DC, pp. 1-14. Jordan, T.H., 1981. Continents as a chemical boundary layer. Phil. Trans. R. Soc. London. A 301, 359-373. Jordan, T.H., 1988. Structure and formation of the continental tectosphere. J. Petrol., Spec. Lithosphere Issue, 11-37. Lenardic, A., 1997. On the heat flow variation from Archean cratons to Proterozoic mobile belts. J. Geophys. Res. 102, 709-721. McKenzie, D.P., Roberts, J.M., Weiss, N.O., 1974. Convection in the earth's mantle: towards a numerical simulation. J. Fluid Mech. 62, 465-538. Morgan, W.J., 1983. Hotspot tracks and the early rifting of the Atlantic. Tectonophysics 94, 123-138. Morgan, P., 1984. The thermal structure and thermal evolution of the continental lithosphere. Phys. Chem. Earth 15, 107-193. Morgan, P., 1985. Crustal radiogenic heat production and the selective survival of ancient continental crust. J. Geophys. Res. 90, C561-C570. Morgan, P., Sass, J.H., 1984. Thermal regime of the continental lithosphere. J. Geodyn. 1, 143-166. Nyblade, A.A., Pollack, H.N., 1993a. A global analysis of heat flow from Precambrian terrains: implications for the thermal
91
structure of Archean and Proterozoic lithosphere. J. Geophys. Res. 98, 12207-12218. Nyblade, A.A., Pollack, H.N., 1993b. A comparative study of parameterized and full thermal convection in the interpretation of heat flow from cratons and mobile belts. Geophys. J. Int. 113, 747-751. Pollack, H.N., 1986. Cratonization and thermal evolution of the mantle. Earth Planet. Sci. Lett. 80, 175-182. Pollack, H.N., Hurter, S., Johnson, J.R., 1993. Heat flow from the Earth's interior: analysis of the global data set. Rev. Geophys. 31,267-280. Price, R.A., 1981. The cordilleran thrust and fold belt in the southern Canadian Rocky Mountains. In: McClay, K.R., Price, N.J. (Eds.), Thrust and Nappe Tectonics, Geol. Soc. London Spec. Publ., Geol. Soc. London, London, Vol. 9, pp. 427-448. Rao, R.U.M., Rao, G.V., Reddy, G.K., 1982. Age dependence of continental heat flow fantasy and facts. Earth Planet. Sci. Lett. 59, 288-302. Rudnick, R., Nyblade, A.A., 1999. The thickness and heat production of Archean lithosphere: constraints from xenolith thermobarometry and surface heat flow. In: Fei, Y., Bertka, C.B., Mysen, B.O. (Eds.), Mantle Petrology: Field Observations and High-Pressure Experimentation. Special pub. in honor of F.R. Boyd, Geochemical Society (in press). Rudnick, R., McDonough, W.F., O'Connell, R.J., 1998. Thermal structure, thickness and composition of continental lithosphere. Chem. Geol. 145, 399-415. Schilling, J.-G., 1991. Fluxes and excess temperatures of mantle plumes inferred from their interaction with migrating mid-ocean ridges. Nature 52, 397-403. Sclater, J.G., Jaupart, C., Galson, D., 1980. The heat flow through oceanic and continental crust and the heat loss of the Earth. Rev. Geophys. Space Phys. 18, 269-311. Smith, A.G., 1981. Subduction and coeval thrust belts, with particular reference to North America. In: McClay, K.R., Price, N.J. (Eds.), Thrust and Nappe Tectonics. Geol. Soc. Spec. Pub., Geol. Soc. London, London, Vol. 9, pp. 111-124. Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution, Blackwell, Oxford, 312 pp. Unesco, Geological World Atlas, International Geological Mapping Bureau, Unesco, Paris, 1976. van der Lee, S., Nolet, G., 1997. Upper mantle S velocity structure of North America. J. Geophys. Res. 102, 2281522838. Vitorello, I., Pollack, H.N., 1980. On the variation of continental heat flow with age and the thermal evolution of continents. J. Geophys. Res. 85, 983-996. Wheeler, J.O., 1987. Tectonic assemblage map of the Canadian Cordillera and adjacent parts of the United States of America. Geological Survey of Canada, Open-file report 1565.
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LITHOS ELSEVIER
@
Lithos 48 (1999) 93-114
The thermal structure and thickness of continental roots C. Jaupart ~'*, J.C. Mareschal
b
a Institut de Physique du Globe de Paris, 4, Place Jussieu, F-75252 Paris Cedex 5, France b GEOTOP, Centre de Recherche en G~ochimie isotopique et en Gdochronologie, Universitd du Qudbec h Montrdal, P.O. 8888, Sta. 'Centre Ville',Montr~al, Canada H3C3P8
Received 6 May 1998; received in revised form 20 July 1998; accepted 10 August 1998
Abstract We compare heat flow data from the Precambrian shields in North America and in South Africa. We also review data available in other less well-sampled Shield regions. Variations in crustal heat production account for most of the variability of the heat flow. Because of this variability, it is difficult to define a single average crustal model representative of a whole tectonic province. The average heat flow values of different Archean provinces in Canada, South Africa, Australia and India differ by significant amounts. This is also true for Proterozoic provinces. For example, the heat flow is significantly higher in the Proterozoic Namaqua-Natal Belt of South Africa than in the Grenville Province of the Canadian Shield (61 vs. 41 mW m - 2 on average). These observations indicate that it is not possible to define single value of the average heat flow for all provinces of the same crustal age. Large amplitude short wavelength variations of the heat flow suggest that most of the difference between Proterozoic and Archean heat flow is of crustal origin. In eastern Canada, there is no good correlation between the local values of heat flow and heat production. In the Archean, Proterozoic and Paleozoic provinces of eastern Canada, heat flow values through rocks with the same heat production are not significantly different. There is therefore no evidence for variations of the mantle heat flow beneath these different provinces. After removing the local crustal heat production from the surface heat flow, the mantle (Moho) heat flow was estimated to be between 10-15 mW m -2 in the Archean, Proterozoic and Paleozoic provinces of eastern Canada. Estimates of the mantle heat flow in the Kaapvaal craton of South Africa may be slightly higher ( = 17 mW m-Z). Large-scale variations of bulk crustal heat production are well-documented in Canada and imply significant differences of deep lithospheric thermal structure. In thick lithosphere, surficial heat flow measurements record a time average of heat production in the lithospheric mantle and are not in equilibrium with the instantaneous heat production. The low mantle heat flow and current estimates of heat production in the lithospheric mantle do not support a mechanical (conductive) lithosphere thinner than 200 km and thicker than 330 km. Temperature anomalies with surrounding oceanic mantle extend to the convective boundary layer below the conductive layer, and hence to depths greater than these estimates. Mechanical and thermal stability of the lithosphere require the mantle part of the lithosphere to be chemically buoyant and depleted in radiogenic elements. Both characteristics are achieved simultaneously by partial melting and melt extraction. 9 1999 Published by Elsevier Science B.V. All fights reserved. K e y w o r d s : Heat flow; Crustal heat production; Mantle heat flow; Lithosphere thickness; Canadian Shield; Precambrian heat
flow
* Corresponding author. E-mail:
[email protected] 0024-4937/99/$ - see front matter 9 1999 Published by Elsevier Science B.V. All fights reserved. PII: S 0 0 2 4 - 4 9 3 7 ( 9 9 ) 0 0 0 2 3 - 7
94
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
1. Introduction Continental heat flow data have been analyzed like the oceanic lithosphere, i.e., within a uniformitarian framework where geological age is the most important factor. The aim was to determine 'average' lithospheric models valid throughout the world (e.g., Polyak and Smirnov, 1968; Vitorello and Pollack, 1980; Morgan, 1985; Nyblade and Pollack, 1993). This procedure was justified in part by the paucity of heat flow data, which required the collection of measurements from different areas to obtain a large enough data set, and in part by the fact that provinces of a given age seemed to have the same average heat flow. However, it is now clear that, even in the oldest provinces of Archean age, the standard deviation of heat flow values (11 mW m -2 for an average of 41 mW m -z, Nyblade and Pollack, 1993) is large and reflects important variations of bulk crustal composition and heat production. The apparent equality of the average heat flow of different provinces may therefore be a statistical artefact and an average heat flow value may not be compared to an average crustal composition determined on the basis of geochemical or petrological models. A long-held belief was that, in Archean and Proterozoic provinces, the crust was on average thinner than in younger provinces and devoid of enriched granitic rocks. This assumption has not been supported by the recent studies of the deep structure of the continental crust (such as conducted by LITHOPROBE in Canada). It has now become apparent that ancient continents are extremely complex assemblages with large variations of crustal thickness and composition (see for instance, Clowes et al., 1992). The main implication for the heat flow studies is the importance of the crustal signal that must be carefully accounted for and makes any kind of generalization difficult. One objective of this paper is to discuss the implications for global heat flow studies of the data that have been collected recently in Canada. As regards lithosphere thickness, we shall emphasize that different surface heat flows lead to different temperature profiles at depth even if the mantle heat flow is kept constant (Pinet et al., 1991). Thus, differences of lithospheric thickness, as deduced from seismological data, for example, do not necessarily provide evidence for mantle heat flow differences.
Because most of the deep mantle nodules studied so far come from Kaapvaal craton and surrounding mobile belts (Nixon et al., 1981), one may ask whether the deep lithospheric structure beneath South Africa is the same as in other cratons. It has been proposed (Jordan, 1981) that continental roots are stabilized by melt extraction, which acts to lower the density and to increase the viscosity of the mantle residue. If this hypothesis is correct, there is no reason to believe that melting has affected the Earth's mantle in the same way everywhere. Thus, depending on the extent of melting, different geological provinces may be underlain by different lithospheres. A second objective of this paper is to use the heat flow data from the Canadian and South-African Shield to compare their deep lithospheric structures. The paper is mostly a review summarizing recent results on heat flow and crustal heat production, focusing on the methods and uncertainties involved in the interpretation of heat flow measurements. It includes a new heat flow map of the Canadian Shield and a new discussion of radiogenic heat production in the mantle part of the lithosphere. For the purposes of using heat flow data for comparison with independent constraints, it is useful to understand the various sources of uncertainties in the measurements and we briefly recapitulate these in Appendix A.
2. Heat flow and heat production in the Canadian Shield Fig. 1 shows the main tectonic provinces of Canada and their extensions into the United States of America. These provinces have different petrological and geological characteristics, and are separated by tectonic boundaries. We have established a heat flow map of the Canadian Shield (Fig. 2), based on all previously published determinations (Jessop et al., 1984, and references therein; Drury, 1985; Drury and Taylor, 1987; Drury et al., 1987; Hart et al., 1994) and on the data that we collected (Mareschal et al., 1989; Pinet et al., 1991; Guillou et al., 1994; Guillou-Frottier et al., 1995, 1996). The Canadian data set is of very good quality because, at many sites, heat flow has been determined using several deep neighbouring boreholes (Appendix A). The data are
95
C. Jaupart, J.C. Mareschal/Lithos 48 (1999) 93-114
~
SUPERIOR = 2.7 Gy 42 _+ 10 m W m -2
0
TRANS-HUDSON = 1.9Gy 42 + 11 m W m -2
i; ~i
jJ
rovinee
J
GRENVILLE = 1.0Gy 41 + 11 m W m -2
APPALACHIANS = 0.4 Gy 57 + 13 m W m -2
Fig. 1. Tectonic map showing the main Precambrian provinces of North America adapted from Hoffman (1989). The Proterozoic Trans-Hudson Orogen is sandwiched between two older Archean cratons. A large data set is available in the Abitibi Belt of the Superior Province. The 'K' region is the Kapuskasing uplift.
summarized in Table 1. For our present purposes, several features of the map of heat flow must be emphasized. (1) There is no apparent change in surface heat flow across the Grenville Front which separates the Archean Superior Province and Proterozoic provinces. The heat flow is uniformly low throughout most of the Grenville Province. A few high heat flow values are scattered in the Central Metasedimentary Belt. These short wavelengths heat flow highs have shallow crustal sources and are related to late intrusives.
(2) Heat flow variations within a given province are as large as between different provinces. The straight average and standard deviation of the heat flow values in the Superior Province is 42 _+ 10 mW m -2. Different subprovinces of the Superior differ by their heat flow fields. The Abitibi Belt has an average heat flow of 37 mW m -z, which is significantly smaller than the average for the rest of the Superior Province, 45 mW m -2 (Table 1). This difference is due to the large amounts of greenstone terranes with low radioactivity in the Abitibi. In the Abitibi, heat flow is lowest ( - - 2 9 mW m -z) near
Fig. 2. Heat flow map of the Canadian Shield. The dots represent the heat flow sites, including the sites in Lake Superior from Hart et al. (1994).
97
C. Jaupart, J.C. Mareschal/Lithos 48 (1999) 93-114
Table 1 Heat flow statistics for the Canadian Shield and Appalachians with the number of heat flow sites, mean heat flow and standard deviation Refs." (1) Pinet et al. (1991) and references therein; (2) Guillou-Frottier et al. (1995) and references therein; (3) Guillou et al. (1994); (4) Drury et al. (1987); (5) Guillou-Frottier et al. (1996); (6) Mareschal et al. (1999); (7) Birch et al. (1968); (8) Jaupart et al. (1982). Province
< Q> •
NO
Superior (exc. Abitibi) Huronian Abitibi Abitibi a Total Superior (Archean) Trans-Hudson Orogen Grenville Grenville a Total Proterozoic Total Precambrian Shield Canadian Appalachians U.S. Appalachians Total Appalachians
26 5 26
45 51 37 39 42 42 41 39 43 41 56 58 57
57 31 30 50 107 43 36 79
(mW m -2)
+ 12 + 7 + 7 + 10 _+ 11 _+ 11
< A > (ixW m -3)
NA
Ref.
21
1.4 + 1.2
21
0.41 + 0.33
(1), (3), (5)
44 29
0.95 _+ 1 0.6 _+ 0.4
(1), (2) (4), (5), (6) (1), (2)
0.8 _+ 11 + 10 + 12 _+ 13 _+ 13
14 36 50
2.6 _+ 2 2.5 _+ 1.9 2.6 +_ 1.9
(1), (2), (4) (7) (1), (2), (4), (8)
aArea weighted average.
the Grenville Front and highest ( = 60 mW m -2) toward the western edge of the province. This long-
wavelength heat flow variation reflects large-scale changes of bulk crustal structure and composition. 58 ~ 4 5 ' N
r~
Athabasca basin
0
WB
HEARNE CRATON
Lynn
Lake Belt
Kisseynew Gneiss Domain - Snow
100 I
km
Saskatchewan I Manitoba 52 ~ 5 0 ' N Fig. 3. Main tectonic belts of the Proterozoic Trans-Hudson Orogen. Four belts (La Ronge, Lynn Lake, Flin Flon, Snow Lake) are made of volcanic sequences and gneisses interpreted to be ancient island arcs. The Thompson Belt is a mixture of reworked Archean basement and continental margin deposits. The Kisseynew and Glennie Domains are made of marginal basin and island arc sequences. WD, WB and RD stand for the Wollaston fold belt, the Wathaman batholith (a continental volcanic and plutonic arc) and the Rottenstone Domain (marginal basin deposits).
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
98
(3) The Paleo-Proterozoic Trans-Hudson Orogen in central Canada is a complicated assemblage of belts with different compositions (Fig. 3). There are marked differences between these various belts. The highest heat flow values are found in the Thompson Belt, on the margin of the Superior craton, and low heat flow values ( < 40 mW m - 2 ) have been recorded in the Lynn Lake Belt in the center of the Reindeer tectonic zone, within the orogen (Guillou-Frottier et al., 1996. Within a belt and from one belt to the other, heat flow varies over short distances ( < 50 km). (4) The lowest heat flow values recorded in the eastern part of the Canadian Shield are 27-28 mW m -2. These values are found in the eastern part of the Abitibi subprovince and in the neighboring part of the Grenville. In the western part of the Shield, low values have been reported at Lynn Lake (Drury, 1985). Without this isolated anomalous site, the lowest values of the Lynn Lake Belt are the same as in eastern Canada. (5) There is a marked increase in surface heat flow in the Appalachians. The boundary between low and high follows almost exactly the limit between the Appalachians and the Grenville, including the Adirondacks. Between the Adirondacks and the Appalachians, the contrast is sharp and occurs over a distance of about 70 km. (6) In many cases, heat flow changes at the boundary between two geological provinces, suggesting that the change is due to different amounts of crustal heat production. The wavelength of heat flow variations provides an independent confirmation. For wavelength A, the ratio between the magnitude of heat flow variations at the surface, Q(0) at depth h, Q(h), is:
Q(O)
1 (1)
Q(h)
cosh(27rA) "
A useful rule of thumb is that this ratio is less than 0.1 for all wavelengths less than = 2.8 h. For a crustal thickness of 40 kin, this corresponds to a threshold wavelength of 110 km. Heat flow variations in the Canadian Shield, which are observed
over wavelengths smaller than this, have a crustal origin.
3. The South African Shield The architecture of the Precambrian Shield of South Africa is similar to that of the Canadian Shield (Fig. 4). The Archean Kaapvaal craton is surrounded by Proterozoic mobile belts: the Namaqua-Natal (1110 Ma) to the South and the Limpopo (2300 Ma) belts. We shall compare the distribution of heat flow values in the Kaapvaal craton and surrounding mobile belts with that of the Canadian Shield. The mean heat flow and heat production within the different Shield provinces of South Africa and Canada are compared in Table 2. (1) Within the Kaapval craton itself, the heat flow is markedly higher in the Witwatersrand basin (51 _ 6 mW m -z) than in neighbouring granitic terranes (33 _+ 2 mW m -z) (Ballard et al., 1987; Jones, 1988, 1992). These differences can be attributed to different crustal heat production values (Jones, 1988). In particular, the thick sedimentary sequence of the Witwatersrand contributes about 7 mW m -2 to the surface heat flow (Jones, 1988). This shows that a straight heat flow average over the whole craton is not a meaningful number because it lumps together different types of crust. Furthermore, the crustal thickness in the craton varies from about 36 km in the central and southern parts (Durrheim and Green, 1992) to about 44 km in the northern part (Qiu et al., 1996). After removing the heat production of the sediments, the mean heat flow of the Witwatersrand basin is only slightly higher than the mean of the Superior Province in Canada. However, this apparent similarity is achieved for different thicknesses of basement rocks: ca. 30 km and 42 km for South Africa and Canada, respectively. We also note that the lowest heat flow values in the Shield of South Africa are 5 mW m -2 higher than the lowest values in the Canadian Shield. (2) An extensive data set for U, Th and K concentrations in upper and middle crustal rocks of the Kaapvaal craton is available across the Vredefort structure (Hart et al., 1981; Nicolaysen et al., 1981).
99
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114 15 ~
~>C..::'.I:..~:-.:3.::""::"::--::"::":i"::"::--
,:ii::i!: ...... 25*
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:-'-
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::: ::'::" '""
=""
++§
+
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9 §
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,
§
.
:...,,,§247247247247247247247247 § § § , , § :§247247 ,//~/
::::::.:: :,, . . . . . . . §
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:: :: :: :: ::" :: :----
§
§
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....................%...
55*
,
I. . . . . . .
I . . . . . . .
15~
20 ~
1. . . . . . . 25 ~
1.
.
.
.
.
50 ~
Fig. 4. Main tectonic provinces of the South African Shield. At the southern edge of the Kaapvaal craton, heat flow increases over a small horizontal distance in the Lesotho region.
A small sample of granulite facies rocks and good geophysical control allow the construction of the full
crustal column beneath the Witwatersrand basin and yield a value of 34 m W m - 2 for the total amount of
Table 2 Comparison of heat flow and heat production in the Archean and Proterozoic provinces of Canada and South Africa, including heat flow (mean and standard deviation, range, number of data) and heat production (mean and standard deviation, number of data) Refs.: (1) This paper; (2) Ballard et al. (1987); (3) Jones (1988); (4) Jones (1992); (5) Jones (1987). < Q>
Range
NO
Na
Ref.
A rchean Superior(Canada) Witwatersrand (S. Africa) Kaapvaal granite domes (S. Africa) Lesotho (S. Africa)
42 51 33 61
10 6 2 16
27-59 36-72 31-36 41-83
57 81 7 9
1.4 + 1.2
44
2.3 ___0.8
7
(1) (2)-(3) (2)-(3) (4)
Proterozoic Trans Hudson (Canada) Grenville (Canada) Namaqua (S. Africa)
42 + 11 41 + 11 61 + 11
23-81 28-68 39-81
31 30 20
0.6 __+0.4 0.8 2.3
29
_ + + +
10
(5)
100
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
heat generated in the crust there. For the mean heat flow in the area (51 mW m -2 ), this leads to a mantle heat flow of about 17 mW m-2 (Jones, 1988). This is ~ 5 mW m -2 higher than the value inferred for the Canadian Shield by Pinet et al. (1991) and Guillou et al. (1994). (3) The Proterozoic Namaqua-Natal Belt has an average heat flow of 61 _+ 11 mW m -2 and a crustal thickness of about 43 km (Green and Durrheim, 1990). The mean heat flow in the Proterozoic mobile belts of South Africa is significantly higher than in the Trans-Hudson Orogen (42 mW m -z) or the Grenville Province (41 mW m -z) of the Canadian Shield. (4) Another distinctive feature of the South African Shield is the rather sharp heat flow variation across the boundary between the Kaapvaal craton and the Natal Belt to the South (Jones, 1992). Heat flow increases from 'normal' Precambrian values of about 45 mW m -2 to values of about 80 mW m -2 toward the southern edge of the craton. This change occurs over a small distance of about 30 km, which indicates that it is mostly due to an increase of heat production in the upper crust (Jones, 1992). This effect is more marked than a similar transition across the Thompson Belt on the western margin of the Superior craton in Canada. Thus, the high heat flows of mobile belts at the edges of cratons do not provide unambiguous evidence for high mantle heat flows. (5) The comparison between the South African and Canadian heat flow data is complicated because the heat flow data from South Africa have not been corrected by Pleistocene and Holocene surface temperature variations (Ballard et al., 1987; Jones, 1987). Such a correction is routinely performed over the areas of Canada that were covered by glaciers. This correction increases the heat flow value by - - 3 - 5 mW m -2. It has now become clear that the ground surface temperature was lower during the last glacial episode even in regions that were not glaciated, although the amplitude of the cooling is not well known for the southern hemisphere. Applying a correction, as has been done in Canada, would increase slightly the heat flow estimates in South Africa, and hence the difference with Canada. One must therefore conclude that, in the African Shield, the mantle heat flow could be slightly larger, and the lithosphere slightly thinner, than in the Canadian Shield.
4. The interpretation of continental heat flow data Most of the studies in continental heat flow were aimed at demonstrating global trends in the data independently of local geological structure. 9 Following Polyak and Smirnov (1968), several workers have tried to find a dependence between heat flow and geological age. This relationship could reflect the relaxation time of the lithosphere following a large-scale orogenic or magmatic event a n d / o r systematic variations in composition with crustal age. 9 Because radiogenic heat production in crustal rocks is large, efforts were made to estimate the heat flow at the Moho (the 'mantle' heat flow). The standard approach has been to compare the local values of heat flow and heat production. In many provinces, within the relatively large errors in both heat flow and heat production values, this relationship seems to be linear. Extrapolating the heat flow data to zero surface heat production yields a 'reduced' heat flow. This relationship was used to define 'heat flow provinces' that have identical characteristics below the surficial heat-producing layer. The relationship between heat flow and age could then be recast as the relationship between reduced heat flow and age. 9 A third method has been to look for systematic geographical trends in the distribution of heat flow values in a continent. Because it was commonly assumed that crustal age decreases from the center to the edge of the continents, this approach was equivalent to the previous ones. At the edges of a continent, heat flow values must grade into oceanic values which are systematically higher. By and large, these interpretations were made at a time when knowledge of crustal composition was inadequate. Indeed, the heat flow data themselves were used to constrain the average crustal composition (Taylor and McLennan, 1985). It should be noted that several studies that used other geophysical constraints to build local crustal models led to estimates of the crustal heat production higher than those suggested by the global studies (e.g., Smithson and Decker, 1974; Smithson and Ramberg, 1979; Pinet and Jaupart, 1987). The available heat flow determinations in the Canadian Shield cover several different geological provinces and allow an analysis of the relationship
101
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
between heat flow, crustal composition and structure. We shall thus use the Canadian data set to discuss these three methods. 4.1. Heat flow vs. age
In the Canadian Shield, the average Archean heat flow, based on a total of 57 individual determinations, is 42 +_ 10 mW m -2. This is almost identical to the result of Nyblade and Pollack (1993) (41 +_ 11 mW m -z) for a worldwide Archean data set which does not include the recent Canadian measurements. The Early Proterozoic Trans-Hudson Orogen (formerly Churchill Province) has an average heat flow of 42 mW m -2 identical to the Archean value. In the Late Proterozoic Grenville Province, the average heat flow is the same as that of the Archean. There is therefore no systematic heat flow variation through the Precambrian of Canada provinces. A marked increase of heat flow occurs in the young Appalachian Province. However, when sites on late and enriched intrusives are excluded, the average Appalachian heat flow is identical to the average values in the Grenville and Superior (Jaupart et al., 1982). This shows that the high Appalachian heat flow is due to a relatively thin radiogenic upper crustal layer. There is therefore no evidence for an enhanced mantle heat flow beneath the Appalachians.
of the Kaapvaal craton in Lesotho, the Thompson Belt at the margin of the Superior, the Adirondacks-Appalachian transition), the lengthscales of heat flow variations are small, demonstrating that changes are of crustal origin.
(a) r
E
3: g 60 ''
N 40~D zZ
~
20
i!I "
9 outside Abitibi 9 within Abitibi
0
1
3 4 Heat production (pWm-3)
5
Grenville
eq
E ~: 60 E
ZZ
40 r
4.2. Heat flow variations at the edge of cratons
Ballard et al. (1987) and Nyblade and Pollack (1993) have suggested that heat flow increases systematically at the edges of cratons. There is no doubt that this is generally true, notably in South Africa, but several explanations can be put forward. Ballard et al. (1987) have suggested that thick and cold Archean continental roots act to divert heat toward the surrounding mobile belts. An alternative explanation is that the younger belts are made of more radiogenic rocks. The Canadian Shield data exhibit more complexity. In the Abitibi, heat flow increases inward, from the craton edge to the center, only to decrease sharply over the Kapuskasing discontinuity. On the scale of the whole Canadian Shield, heat flow highs are recorded in the West Abitibi and in the Thompson Belt, and there is no clear geographic trend across the Shield. In several areas (the margin
Superior province
2oi o
i
2
3
4
Heat production (gWm-3) (C)
Trans Hudson Orogen 100|
~,
9 Flin F l o n S n o w L a k e Belt | 9 LynnLake Belt
801
9 T h o m p s o n Belt
./ o
9
9
.
.
*
..a (D
7~ 20 0
2.0 HEAT PRODUCTION (pWm-3)
Fig. 5. Heat flow and heat production data for the different provinces of the Canadian Shield. (a) Superior Province (age > 2500 Ma); (b) Grenville Province (age ca. 1100 Ma) and (c) Trans-Hudson Orogen (age ca. 1800 Ma).
102
c. Jaupart, J.C. Mareschal/ Lithos 48 (1999) 93-114
Fig. 5 shows the heat flow and heat production data from the various provinces of the Canadian Shield. The correlation is obviously very weak within a single province, such as the Grenville (Fig. 5). Selecting data from a subprovince with given geological evolution does not lead to a marked improvement. In the Abitibi, for example (Fig. 5), the poor correlation between heat flow and heat production is readily explained by the presence of a surficial layer of greenstone rocks that are depleted in radioactive elements (Pinet et al., 1991). Surprisingly, the values of the parameters qr and D that are obtained by linear regression to all the available data are similar for the Superior, the Grenville, and the Appalachian Province (Table 3). Given the errors on the values of these two parameters (-~ 2 mW m -2 and 2 km, respectively), one may not distinguish between these three provinces. In the Trans-Hudson Orogen (Fig. 3), the various belts exhibit different trends of heat flow vs. heat production (Fig. 5c). In the Flin FlonSnow Lake Belt, the heat flow is almost constant regardless of the surface heat production (D ~ 0); the correlation is better for the Lynn Lake Belt but the slope, D, of the apparent heat flow-heat production relationship is unusually large (26 km vs. a standard value of 8-10 km); the correlation is also poor in the Thompson Belt (Guillou-Frottier et al., 1996). The parameters of the heat flow-heat production relationship for all the sites in the Trans-Hudson Orogen (qr = 33 _+ 3 mW m -2 and D = 14 _+ 4 km) suggest a poorly differentiated crust. The heat flow-heat production relationship was initially established over exposed plutons very enriched in radioactive elements (Birch et al., 1968). If their contribution dominates the heat flow signal, it can be inferred from the heat flow-heat production relationship. These conditions are not encountered in
4.3. The relationship between heat flow and heat production
Birch et al. (1968) had proposed that, within a heat flow province, there is a linear relationship (2)
q=qr+aD
between the observe heat flow q and A the local heat production of rocks of the crystalline basement. The slope D, which has dimension of length, is related to the thickness of a surficial heat producing layer and qr is called the reduced heat flow. Among the many heat sources distributions that fit the linear relationship, the exponentially decreasing was favored because it is independent of the erosion level (Lachenbruch, 1970). For the values of D (commonly = 10 km), this would imply that the total heat production in the crust is AD and that the mantle heat flow = qr (In early papers, the reduced heat flow was noted as qm). Most of these conclusions and the linear relationship itself have been challenged by the recent studies. Measurements of heat production in exposed crustal sections such as the Vredefort in South Africa (Nicolaysen et al., 1981) or the Pikwitonei-Sachigo and Kapuskasing-Wawa areas of Canada (Ashwal et al., 1987; Fountain et al., 1987) have shown that the concentration of heat sources does not decrease with depth as would be required by the heat flow-heat production relationship. On a vertical scale of -- 10 km, measurements on samples from the deep boreholes at Kola, Russia, (Arshavsakaya et al., 1987) and the KTB, Germany, (Clauser et al., 1997) indicate no systematic variation of heat production with depth. Finally, as the number of data increased, it appeared that the correlation between surface heat flow and heat production is weak. This is very well demonstrated by the Canadian heat flow data.
Table 3 Heat flow heat generation relationship in Northeastern America Province < Q > (mW m-2)
Qr (mW m-2)
D (km)
# Values
Superior (inc. Abitibi) Grenville Trans-Hudson Orogen Canadian Appalachians U.S. Appalachians
28 30 33 40 33
10 9 14 10 8
85 30 29 14 36
40 41 42 59 57
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
the Canadian Shield and other approaches must be used to constrain the crustal heat production. The same observations were made in South Africa (Jones, 1987, 1988).
5. Other Shield areas
The discussion has so far been focused on two specific continental regions and it is worthwhile to discuss briefly other cases. For our present purposes, we restrict ourselves to those areas where a minimum number of heat flow and heat production measurements are available. 5.1. Australia The Australian continent can be split into three main regions. The Western Shield region is made of two Archean cratons (Yilgarn and Pilbara) and associated greenstone belts. The Central Shield region has a large Proterozoic craton and mobile belts. The eastern part of Australia has been geologically active recently. Table 4 summarizes the data available for the Archean and Proterozoic Shields. Three facts strengthen the conclusions reached above. The average heat flow through the Yilgarn craton is 37 mW m -z , which is smaller than the average heat flow through the whole Western Shield (49 mW m-Z). This is reminiscent of the Canadian case, where the Abitibi Province differs from the Canadian Archean as a whole. A second fact is that the average Archean heat flow is smaller than the average Proterozoic heat flow of 78 mW m - 2 (Table 4). This difference parallels that of the average heat production of sur-
Table 4 Heat flow and heat generation in Archean and Proterozoic provinces of Australia Refs.: (1) Sass and Lachenbruch (1979); (2) Cull (1991).
NQ
< A >
Na
Ref.
Archean Yilgarn craton Total Western Shield
37 _+6 49_+ 13
15 37
2.6_+2.1
12
(1) (2)
Proterozoic Gawler craton Total Central Shield
82 + 24 7 8 _ 19
6 75
3.6+ 1.9
38
(1) (2)
103
face rocks and hence does not provide evidence for mantle heat flow differences beneath the two Shield regions. Unfortunately, there are no detailed models of crustal composition for both regions which would allow estimates of the mantle heat flow. However, from an analysis of the relationship between heat flow and surface heat production, Sass and Lachenbruch (1979) found that the two Shield regions are characterized by the same 'reduced' heat flow and hence concluded that differences are limited to the upper crust. The third fact is that heat flow in Australia does not conform to a simple concentric distribution away from the center of the continent: the Central Shield has higher heat flows that the Western Shield at the edge. Taken as a whole, the Australian data demonstrate that heat flow differences are accounted for by differences of crustal heat production and follow the geological assemblage. 5.2. Indian Shield A significant number of heat flow measurements are available in the Southern Indian Shield (Gupta et al., 1991, 1993). The Archean Dharwar craton can be divided in two geologically distinct regions in the East and in the West. The former has numerous granitic bodies intruding gneisses and the latter is mostly made of supracrustal granite-greenstone belts. This difference is reflected in the values of heat flow, which average to 40 _+ 3 mW m -2 and 31 _+ 4 mW m - 2 in the Eastern and Western Dharwar cratons, respectively (Gupta et al., 1991). There again, one sees that the Archean continental crust is a heterogeneous assemblage whose composition varies from one region to the next. The heat flow differences between the two parts of the craton are due to different rates of crustal heat production and do not imply deep-seated thermal contrasts. The Proterozoic Bastar craton in the central Indian Shield has larger heat flow values than its Archean counterparts. Three reliable measurements range from 51 to 64 mW m -2 (Gupta et al., 1993). These heat flow values are found through rocks with larger values of radiogenic heat production than in the Archean. For this reason, Gupta et al. (1993) conclude that the contrast in surface heat flow between the Archean and Proterozoic cratons of India is due to different upper crustal compositions.
C. Jaupart, J. C. Mareschal / Lithos 48 (1999) 93-114
104
In India, therefore, heat flow is mostly sensitive to crustal composition and its distribution does not conform to any simple geometric pattern.
5.3. East China Shield East China has been subjected to recent thermal events in the Late Mesozoic and Cenozoic implying that measured heat flow values are not in secular equilibrium with crustal heat production and mantle heat flow. Gao et al. (1998a,b) have reported a large data set for U, Th and K concentrations in rocks of all types and all metamorphic grades from the East China Shield. They present a series of detailed crustal models for various regions of East China and derive estimates of the bulk crustal heat production. One important conclusion is that crustal structure and thickness vary laterally by large amounts, implying that a single thermal model cannot account for the geological complexity of East China.
6. Mantle provinces
heat
flow
beneath
Precambrian
The detailed studies in Canada and South Africa demonstrate that large heat flow differences exist within a single province, related to differences of crustal structure. Crustal columns in different areas may vary in thickness and in composition. For example, the western part of the Abitibi Belt of Canada has a large thickness of granodioritic/tonalitic rocks and a shallow greenstone cover, and heat flows as large as 59 mW m -2. In contrast, the eastern part of the Abitibi has a thick greenstone sequence and fewer granodioritic/tonalitic plutons, which is reflected in small heat flows of about 29 mW m -2. In the Kapuskasing uplift, granulite facies terranes have been brought to the surface by low-angle thrusts and make the bulk of the crustal column. A heat flow average through this area of the Canadian Shield therefore mixes different types of crustal structure and rocks, and does not correspond to any well-defined petrological crust. Similar conclusions have been reached for the other Precambrian shields.
6.1. The crustal contribution to the surface heat flow The amount of mafic rocks in the deep crust is a key variable because it affects the crustal heat production estimate and reflects the extent of past melting events in the mantle. It has a strong effect on the average seismic velocity of the lower crust and hence may be estimated with seismological techniques (Christensen and Mooney, 1995; Rudnick and Fountain, 1995). For the purposes of estimating the crustal contribution to the surface heat flow, unfortunately, such knowledge is not sufficient. The lower crust may contain substantial amounts of felsic and enriched rocks, such as granulite-facies metapelites, for example, which have highly variable trace-element compositions and hence highly variable heat production rates (Rudnick and Fountain, 1995). Thus, it is not possible to relate heat production to seismic velocity (Fountain, 1986), and it is essential to determine the local values of heat production of each crustal structural level in each province. These problems have been discussed at length in Jaupart et al. (1998). Here, we briefly recapitulate some recent data which emphasize that the lower crust is not depleted in radioactive elements and contributes a significant amount of heat. Gao et al. (1998a,b) have derived a series of petrological and seismological models for the crust in the East China Shield. There, the lowermost crust is slightly more evolved, and slightly more radiogenic, than the average Precambrian crustal models of Rudnick et al. (1998) and Christensen and Mooney (1995). In the East China Shield, for an average thickness of 37 km, the total amount of crustal heat is 37 mW m -2, i.e., the average crustal heat production is 1.0 txW m -3. In the Indian Shield models of Gupta et al. (1991), the total Archean crustal heat flow contribution is between 26 and 21 mW m -2 for 38-km thick crust. This constrains the average Archean crustal heat production in India to be between 0.55 and 0.68 txW m -3. On average, the Archean crust beneath the Witwatersrand in South Africa has a heat production rate of 0.94 txW m -3. These independent analyses suggest that there is no such thing as a single representative Archean crustal composition. They further suggest that differences of crustal composition may account for observed heat flow variations.
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
The Witwatersrand heat flow and a crustal model based on exposures in the Vredefort structure lead to a mantle heat flow value of 17 mW m -2 for the Kaapvaal craton (Jones, 1988). With a similar procedure, Gupta et al. (1991) arrived at a range of 12-17 mW m -2 for the mantle heat flow in the Indian Shield. The difference between these estimates, = 5 mW m-2, is within the expected error range for heat flow and bulk crustal heat production.
6.2. Local studies in the Canadian Shield The mantle heat flow is a small residual value, and hence is sensitive to even small errors on the crustal heat production estimates. In the Canadian Shield, U, Th and K concentrations have been determined in all types of crustal rocks. In addition, various crustal levels can be sampled allowing reconstruction of crustal columns below and around measurement sites. The upper and middle crust have been studied over vast areas through systematic sampling (Eade and Fahrig, 1971; Shaw et al., 1986). Using geological maps, data from Ashwal et al. (1987), Fountain et al. (1987), as well as many other published reports, Pinet and Jaupart (1987) and Pinet et al. (1991) have determined the areas for each rock-type and calculated an area-weighted average in two well-sampled different granulite-facies terrains (the Kapuskasing and Pikwitonei areas) and three different amphibolite-facies terrains from the Superior Province. For the granulites, they found the same average heat production value of 0.40 txW m -3. For the three amphibolite-facies terrains, average heat production values are between 1.0 and 1.2 txW m -3. These different data sets are consistent, and come from regions directly relevant to the available heat flow data. Pinet et al. (1991) and Guillou et al. (1994, Guillou-Frottier et al. (1995) have looked in great detail at the relationship between heat flow and local crustal structure in the Archean Abitibi subprovince and its boundaries with the Kapuskasing structural zone. Heat flow increases systematically from the Grenville front to the Kapuskasing uplift over a distance of = 500 km. There is also a sharp drop of about 20 mW m -2 over a distance of 70 km across the Ivanhoe Lake fault, which separates the Abitibi from the Kapuskasing uplift. This demonstrates that
105
these variations of heat flow are of crustal origin (Fig. 2). The long wavelength variation in heat flow is accompanied by an increase in Bouguer gravity. Taken together with seismic results, these two data sets provide constraints that restrict crustal models and the mantle heat flow to a very narrow range. In the Abitibi subprovince, three crustal lithologies dominate: greenstones, tonalite-trondjemite-granodiorite (TTG) and granulite facies rocks. Large outcrops of these three lithologies are exposed in the Shield, allowing direct measurements of their densities and heat production rates. Guillou et al. (1994) generated a series of crustal models by varying the mantle heat flow, the thicknesses of the three lithological units, their densities and heat production rates. These variables are varied within rather large ranges to allow for local departures from the regional average values. Furthermore, a crustal model is considered acceptable if it generates a heat flow value within 10% of the measured value. With this procedure, data uncertainties are taken into account. The constraints of both gravity and heat flow data can be met only when the mantle heat flow lies between 7 and 15 mW m -2. The most probable value is 13 mV~r m -2"
Independent studies in large areas of low heat flow over depleted crust in the Superior and Grenville lead to estimates between 10 and 13 mW m-2 (Pinet et al., 1991; Mareschal et al., submitted). These arguments rule out values smaller than 10 mW m-2, and lead to a range of 10 to 15 mW m -2 for the mantle heat flow beneath the Canadian Shield.
7. Thermal structure and thickness of the lithosphere
7.1. The base of the lithosphere The base of the lithosphere cannot be defined without a specific model for heat transport processes (Fig. 6). There must be a boundary layer through which heat is brought from the convecting mantle into the lithosphere. Below this convective boundary layer, lateral temperature variations are small. Temperature is a continuous variable and hence global geophysical techniques, such as seismic tomography, are only able to determine the sum of the conductive
106
C. Jaupart, J.C. Mareschal/Lithos 48 (1999) 93-114
0
Tm
Temperature
~.~,
Conductive boundary layer
Boundary layer
Convecting mantle
Fig. 6. Sketch illustrating various definitions of lithospheric thickness, h 1 corresponds to the stable part of the lithosphere, where heat transport is achieved by conduction, h 2 is obtained by downward continuation of temperature to an isentropic temperature profile representative of the fully convecting mantle, h 3 is the total thickness over which temperature contrasts with the surrounding oceanic mantle persist. Large-scale tomographic images allow estimates of h 3.
layer and the boundary layer below it, down to a depth which we shall note h 3 (Fig. 6). In order to make a quantitative comparison between the different thickness estimates, one must specify the structure of the convective boundary layer. One possibility if that this boundary layer is unstable and generates small-scale convective motions (Parsons and McKenzie, 1978; Davaille and Jaupart, 1994; Doin et al., 1997). Another possibility is that it is stable and develops as continents sweep over the convecting mantle (Gurnis, 1988; Lenardic and Kaula, 1995; Lenardic and Kaula, 1996). Questions about the long-term survival of the lithosphere, geochemical reservoirs isolated from the fully convecting mantle and transient thermal effects in the lithosphere, deal with the conductive layer which has thickness h 1 (Fig. 6). This is only part of the total thermal boundary layer, and h 1 < h 3. Finally, one must introduce a third thickness value, h2, c o r r e s p o n d i n g to the value obtained by downward continuation of temperature to an isentropic 'oceanic' profile. As shown in Fig. 6, this leads to an intermediate value of the lithosphere thickness, i.e., one has h 1 < h 2 _< h 3. By requiring that the base of the lithosphere lies along a well-mixed isentropic 'oceanic' tempera-
ture profile, one makes the additional assumption that mantle temperatures below continents are the same than beneath oceans, which may not be valid, as shown by dynamical simulations of the effect of continents at Earth's surface (Gurnis, 1988; GuillouFrottier et al., 1995; Moresi and Lenardic, 1997). This effect, however, is expected to be small. The main point is that the 'cold roots' of seismic tomography must extend deeper than both h~ and h2, i.e., the thicknesses which may be derived from thermal considerations. In the convective boundary layer at the bottom of the lithosphere, temperature tends asymptotically to the fully convective profile. Thus, an additional definition is required to calculate the thickness of this layer. In some calculations, the base of this layer is such that temperature reaches a prescribed percentage of the fully convecting mantle temperature. In another definition, the conduction profile is extrapolated linearly until it meets the convecting mantle temperature. This second procedure is analogous to the one used to calculate h2, and hence the convective boundary layer thickness is equal to the difference between h 2 and h~. Jaupart et al. (1998) have recently investigated the small-scale convection
C. Jaupart, J.C. Mareschal/ Lithos 48 (1999) 93-114
model using likely values of mantle rheological parameters and have found a value of 40 km for this thickness difference.
7.2. Vertical temperature profiles We compare changes of deep lithospheric structure across the Abitibi Province with an average thermal model for the younger Appalachian Province. The western part of the Abitibi is characterized by heat flows that are large by Archean standards (59 mW m -z) and contrasts with the eastern Abitibi which has some of the lowest heat flow values found in the Shield (28 mW m-Z). The Appalachians are associated with a high average heat flow of 57 mW m -2, due to a relatively thin enriched upper crustal layer (Jaupart et al., 1982). We have calculated geotherms for different areas using constraints from heat flow data and models for the distribution of heat production in the crust (Table 5). The crustal model has a maximum of three layers with different heat production values. In most calculations, the deepest layer has a heat production of 0.40 IxW m -3, corresponding to granulite-facies terranes in Canada and in the Norwegian Shield. The latter comes from studies in the Egersund-Bamble area of southern Norway (Pinet and Jaupart, 1987), which was one part of the same geological province than the Grenville and hence provides a window into Grenvillian lower crust. The Shield crustal models have
Table 5 Thermal models for eastern Canada WA1 and WA2 are the West Abitibi models. EA is East Abitibi, and Appl and App2 are the Appalachian models. WA1 Qs (mW m -2 ) db a~ db A~ db a~ Crustal
Tr~ (~
52 41 0.98 -
thickness (km)
41 523
WA2 52 30 1.2 11 0.4 -
41 472
a Temperature at the Moho. bLayer thickness (km). c Heat production in layer (txW m - 3 ).
EA 28 20 0.2 28 0.4 -
48 385
Appl
App2
58 8 3.1 10 1.1 22
58 8 2.5 32 0.8
0.4
40 402
-
40 477
107
been discussed above (Pinet et al., 1991; GuillouFrottier et al., 1995). In the Appalachians, heat flow and gravity data demonstrate that surface heat production rates do not extend deeper than a few kilometres (Jaupart et al., 1982; Jaupart, 1983). We first neglect heat production in the mantle part of the lithosphere, which leads to lower bounds for the lithosphere thickness h 2 (Fig. 6). Over the large horizontal distances which separate the regions of Table 4, lateral heat transfer is inefficient and thermal differences extend over great vertical distances, from the surface to the base of the lithosphere. In western Abitibi, heat flow increases toward the edge of the province, toward the Kapuskasing uplift, with a very large wavelength. However, it drops to a small value over a short wavelength on the uplift itself. One must expect some horizontal heat transport in this area. The highest western Abitibi heat flow is 59 mW m -2 and we use a regional average of 52 mW m-2 for this part of the Superior Province. In the crust, thermal conductivity is fixed at 2.5 W m -1 K -1. For the mantle, we account for both lattice conduction and radiation and take the following equation:
k ( T ) = 0.174 + 0.000265T
+
0.368
10 - 9 T 3
(3)
where T is the absolute temperature. This takes into account measurements from several laboratories (Schatz and Simmons, 1972; Beck et al., 1978; Sch~irmeli, 1979), and gives conductivity values which are slightly smaller than those of Schatz and Simmons (1972) at temperatures below 1200 K. Following our analysis, the Moho heat flow is taken to be 12 mW m -2. The crustal models and the resulting Moho temperatures are summarized in Table 5. One source of uncertainty is the vertical distribution of crustal heat production. Different crustal models for the same total heat production indicate that Moho temperature estimates may be inaccurate by several tens of degrees (Table 5). Moho temperatures are coldest beneath the East Abitibi (385~ and highest (523~ in the West Abitibi area. Moho temperatures beneath the young Appalachian Province may be slightly smaller than those beneath the much older Abitibi. Fig. 7 shows the two extreme geotherms through the
108
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
Temperature (C) ,~-.. \ .. \'.. \ \ \ \ \ \ \ \ \
..~ -125
\ \
\
",. '..
\ \
".. \
'. \
_ _ _ East Abitibi
........ West Abitibi
9 \ \
".. '.. \
".. \
.. \
Isentropic profile
independent data such as (P,T) estimates from mantle nodule compositions or the diamond stability field because of uncertainties in the values of thermal conductivity and heat production in mantle rocks. The former source of uncertainty is due in part to the rather small number of carefully conducted measurements and in part to a problem of a more fundamental nature. In the mantle, crystal dimensions are close to the photon mean free path (Schatz and Simmons, 1972; Nicolas and Poirier, 1976). Scattering at grain boundaries is the dominant effect and limits the efficacy of radiative heat transport. A proper determination of bulk mantle conductivity must therefore take into account crystal shapes, sizes and orientations.
9
7.3. Heat production in a thick lithosphere
\ \
-300 Fig. 7. Highest and lowest geotherms for eastern Canada. Geotherms for the Grenville Province and for the Appalachians would fit between these two extreme values (see Table 5).
whole lithosphere. We arbitrarily define the base of the lithosphere to be at well-mixed convective mantle temperatures, along an isentropic profile with a potential temperature of 1280~ (McKenzie and Bickle, 1988). With these definitions, the lithosphere extends to depths of about 290 and 250 km under the coldest and hottest parts of the Abitibi subprovince. Thus, significant differences of deep lithospheric structure are predicted away from known geological boundaries. These estimates of lithosphere thickness h 2 c o r r e s p o n d to a mantle heat flow of 12 mW m -2, which is our best estimate. However, mantle heat flow values of up to 15 mW m-2 are not ruled out by the data, and would lead to decrease the thickness estimates by about 50 km. In the Kaapvaal craton, the 18 mW m -2 variation in surface heat flow imply a difference in Moho temperature of = 125 K between the Witwatersrand and the neighboring regions. These calculations allow a comparison between different provinces for a given set of assumptions on thermal conditions and thermal properties. However, they must be used with caution for comparisons with
In order to avoid deep temperature extrapolations and uncertainties in the values of thermal conductivity, one may use constraints on the mantle heat flow value and estimates of heat production in the cratonic mantle. If the continental lithosphere is very thick, as advocated by Jordan (1981), for example, the thermal time-constant is very large and one must expect a significant component of heat flow due to transient effects. One transient effect which has not been studied is that of radiogenic heat production. If the mantle part of the lithosphere contains a non-negligible amount of U, Th and K, one must account for the rundown of radioactivity with time. The time-constant of the decay of bulk radiogenic heat production depends on the relative amounts of the four isotopes, 238U and 235Th, 232Uand 4~ which have different half-lives. For 'average' T h / U and K / U ratios, the bulk heat production decreases by a factor of about 3 in 3 billion years. One may approximate the heat production by an exponential function: A ( t ) = A 0 e x p ( - at). The characteristic thermal relaxation time of the lithosphere is r = if~K, where L is the thickness of the lithosphere and K is the thermal diffusivity. Here L corresponds to thickness h~ in Fig. 6. It is shown in Appendix B that the contribution of the lithospheric heat production to the surface heat flow is: tan~a~
AQ(t) =AL
v/aT e x p ( - t~t).
(4)
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
This is compared to the instantaneous amount of heat produced in the lithosphere: AQ0 ( t ) = AL exp( - cet).
(5)
The ratio between these two quantities gives the magnitude of the time-delay effect: 2~Q ( t )
tan f~-r =
(6)
AOo( t)
'
which is an increasing function of lithosphere thickness. The ratio reaches a value of 1.5 for a 300-km thick lithosphere. This shows that surficial heat flow measurements record some time-average of deep heat production and are not in equilibrium with the instantaneous lithospheric heat production. Fig. 8 shows the contribution of radiogenic elements in the lithosphere to the surface heat flow. It is useful to compare this with the heat flow which is conducted through the lithosphere in steady state. For the sake of argument, we consider the simplest case of a lithosphere lying on top of a mell-mixed mantle with a uniform temperature of 1573 K. The conducted heat flow is then
Qc=k
AT (7)
h '
where A T = 1300 K, h stands for the lithosphere thickness and the value of k has been fixed at 3.0 W m -1 K -1. It may be seen that, for all values of
40
_
r
30
'
~
,,
_
!
on
'
/
', lithosphere /f ',,!approximate/)/
20
1/I
'
/
/ /
//
_
<
'
/
10 Z 9 r
0
100
2oo
300
400
LITHOSPHERE THICKNESS (km) Fig. 8. Contribution of heat production in the mantle lithosphere to the present surface heat flow. Numbers along the curves stand for values of heat production (~zW m-3). The dashed curve corresponds to steady-state conduction over a temperature difference of 1300 K, for a thermal conductivity of 3.0 W m-1 K-1
109
mantle heat production A larger than 0.02 IxW m -3, the radioactive contribution is comparable to the conducted heat flux in lithospheres thicker than 200 km (Fig. 8). In a 300-km thick lithosphere, mantle heat production must be smaller than 0.01 IxW m -3 to contribute less than 5 mW m -z. Rudnick et al. (1998) have recently compiled measurements of U, Th and K concentrations in different types of upper mantle samples and have found a large range of heat production rates. The lowest values belong to large fragments of the upper mantle which have been brought to the surface by tectonic processes ('massif' peridotites). Rudnick et al. (1998) favor data from peridotite xenoliths found in alkali basalts, and propose a cratonic mantle heat production of 0.02 txW m -3. For comparison with heat flow constraints, one must take into account heat conduction along an isentropic profile, which contributes a heat flux of about 2 mW m -2. Thus, using our constraints on the mantle heat flow value, the maximum amount of heat produced in the Canadian lithosphere is between 8 and 13 mW m -2. A cratonic heat production of 0.02 IxW m -3 rules out lithosphere thicknesses larger than 330 km. For thicknesses larger than 300 km, the time-constant for thermal diffusion in the lithosphere is comparable to, and may even be larger than, the age of geological provinces in the Canadian Shield. In such conditions, surficial heat flow records an additional transient associated with relaxation of the thermal structure achieved at the end of the last tectonic or magmatic event. This argument shows that very thick lithospheres ( > 400 km) can be ruled out if radiogenic heat production in the mantle part of the lithosphere is as large as suggested by Rudnick et al. (1998), independently of uncertainties in the values of thermal conductivity. Partial melting, required to stabilize a deep and cold lithospheric root (Jordan, 1981; Doin et al., 1997), also leads to the depletion of incompatible elements such as U, Th and K. If the 'massif' peridotites offer representative samples of the cratonic mantle, heat production in the deep lithosphere may be as small as 0.006 IxW m -3 (Rudnick et al., 1998). In this case, heat production does not allow useful constraints on lithosphere thickness and the temperature profiles of Fig. 7 are a reasonable approximation. A key question is whether depletion
110
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
processes have affected the Canadian and South African lithospheres with the same intensity. The difference between the mantle heat flow estimates for the two shields may well be due to differences of cratonic mantle heat production. 7. 4. Discussion
Continents are not immobile at Earth's surface and it is unlikely that thermal conditions at their base can be kept constant for long times. For a stable lithosphere thickness h I equal to 250 km and a thermal diffusivity of 10 -6 m 2 s - l , it may be shown that surface heat flow does not record changes of basal conditions over periods smaller than 1 Ga. Thus, there may be no relationship between the present-day deep mantle circulation and continental heat flow. Heat flow measurements allow constraints on the stable and conductive part of the lithosphere. For the convecting mantle, however, it is the whole thermal boundary layer which matters (the 'thermal' lithosphere) and this cannot be defined without consideration of the dynamical interactions between a thick root and the surrounding mantle. The key point is that continents cannot be treated as passive. In this paper, we have shown that one must allow for thermal differences imposed from the top down: known contrasts in crustal heat production lead to significant differences of deep thermal structure which may play an important role for mantle convection.
8. Conclusions (1) Interpretations of continental heat flow data often refer to the linear heat flow-heat production relationship. In the Canadian Shield, this concept is not very useful. All the data from the different provinces in the eastern Canadian Shield fit equally well (or, to be more accurate, equally poorly) the same heat flow-heat production relationship. (2) Geographic variations of the heat flow are found within Precambrian provinces. These variations reflect changes in crustal heat production. Because these variations are very often systematic and related to large-scale changes of crustal structure, an
average value of heat flow is difficult to compare with a synthetic crustal model derived from petrological or chemical considerations. (3) The heat flow is significantly higher in the Proterozoic provinces of South Africa than in those of North America. The apparent relationship between heat flow and crustal age, obtained from worldwide averaging of the heat flow values from provinces of the same age, does not hold when different provinces of the same continents are compared. (4) In Canada and in South Africa, large amplitude (-- 40 mW m -z ) heat flow variations are found over short distances. Such variations can only have a crustal source. In Canada, there is no detectable variation in mantle heat flow between provinces of the Shield. In South Africa, a difference in mantle heat flow between adjacent Proterozoic and Archean provinces is not required by the data. (5) The estimated values of mantle heat flow (10-15 mW m - 2 in Canada and 17 mW m - 2 in the Kaapvaal craton) yield constraints on lithospheric thickness and heat production in the mantle lithosphere. Lithosphere thicknesses larger than 330 km are inconsistent with heat production measurements in mantle xenoliths from South Africa. Thicknesses smaller than 200 km are not consistent with the mantle heat flow values. (6) Different geophysical interpretations do not rely on the same definition of lithosphere thickness. (7) Deep temperature extrapolations may be inaccurate due to uncertainties in thermal conductivity in the cratonic mantle. Determination of the base of the lithosphere requires knowledge of the mechanism of heat transport into the lithosphere. One may choose to emphasize these difficulties, but a more optimistic conclusion is that, despite these potential problems, it is now possible to focus on a specific range of thickness estimates. One may anticipate that even slight improvement on a single variable will considerably tighten the range.
Appendix A. Accuracy of heat flow determinations Heat flow values are determined in the field, in imperfectly controlled environments. Errors do not arise from intrinsic instrumental problems, but from
C. Jaupart, J.C. Mareschal / Lithos 48 (1999)93-114
geological and climatic noise. Because heat flow and temperature profiles are compared to independent data, it is useful to recapitulate the key steps in the measurement technique and the errors involved. In principle, all problems could be avoided if very deep boreholes (deeper than -- 12000 m) could be used. This is seldom the case and there are several types of difficulties. (1) One problem is that rocks may be permeable at shallow depths and interstitial water may transport a significant amount of heat which goes undetected. At large depths, fractures and most pores close due because of high confining pressure, and permeability tends to very small values. (2) Noise can be introduced by variations of the boundary condition at the surface. Effects of topography and lakes can easily be detected and accounted for. More subtle changes, such as due to the effect of the vegetation on the ground surface temperature, can remain undected. When the temperature gradient is small, as often in the Canadian Shield, these perturbations of the gradient can exceed 10% near the surface. They are more easily detected and eliminatedwhen measurements are available from several boreholes. (3) The third type of problems may arise because of lateral heat transfer, due for example to heat refraction in a borehole located near the boundary between rocks with different values of conductivity. When the borehole extends through a narrow body with a strong contrast in conductivity, the local enhancement of the heat flow can be as large as 50% (Guillou-Frottier et al., 1996). Refraction effects are difficult to account for without complete knowledge of the subsurface structure, and the best procedure is to assess directly the magnitude of lateral heat transfer by temperature measurements. One method is to verify that throughout the depth range of the borehole, the local value of the vertical heat flow at each depth is the same. Refraction effect can also be detected when several nearby boreholes are available. Their common value is then assigned to the site. (4) One of the main difficulties lies in estimating the thermal conductivity. Good quality conductivity measurements are made with the 'divided bar' technique. In order to detect local 'grain' effects due to heterogeneity in rock composition, measurements are
111
made on cylinders of different thicknesses. The bulk thermal conductivity value must not depend on sample thickness. These measurements are time consuming because they must be made in thermal equilibrium. Because of the limited number of core samples that can be measured, it is important to sample properly the lithologies in the borehole and to eliminate unrepresentative samples. Poor sampling can result in ~ 20% error on the estimated heat flow. (5) Climatic variability has caused changes in the boundary condition at Earth's surface. Such perturbations propagate slowly through rock and induce a thermal transient at depth. The magnitude of these transients decreases with depth and eventually becomes negligible. For very recent climatic changes (the past 200 years), temperature perturbations are negligible below ~ 200 m. Pleistocene climatic effects extend over more than 3 km, and most boreholes are shallower than this. Sass and Lachenbruch (1979) analyzed temperature data from a very deep borehole in the Proterozoic Flin Flon Belt and found no evidence for Pleistocene effects. Unfortunately, their study did not provide a definite answer because of uncertainties in the conductivity data. The practice has been to make an adjustment for glacial-interglacial climate variability only on heat flow data from regions that were covered by glaciers, although it is now clear that temperature were affected throughout the northern hemisphere. The standard procedure is thus to make a correction based on a model of Pleistocene climate variations (Jessop, 1971). In the Canadian Shield, heat flow and temperature gradients are low and hence the correction is small, typically 2 - 3 mW m -2. The relative error on the magnitude of this correction may be large, but the resulting error on the heat flow is small and within the range of measurement uncertainty. However, the correction is systematic and increases the heat flow value. Pleistocene glaciation was most prominent in the northern hemisphere and should not affect temperatures in the southern hemisphere. Indeed, uncorrected heat flow values seem to be systematically higher in the southern hemisphere than in the northern hemisphere (Sclater et al., 1981). These differences disappear when the northern hemisphere data are corrected. Heat flow values from South Africa have not been corrected for Pleistocene climatic effects.
112
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Appendix B. Heat sources in the lithosphere
References
In order to calculate the effect on the heat flow of the rundown of the heat sources in the mantle lithosphere, we solve the heat equation:
Arshavsakaya, N.I., Galdin, N.E., Karus, E.W., Kuznetsov, O.L., Lubimova, E.A., Milanovski, S.Y., Nartikoev, V.D., Semaskko, S.A., Smirnova, E.V., 1987. Geothermic investigations. In: Kozlovsky, Y.A. (Ed.), The Superdeep Well of the Kola Peninsula. Springer-Verlag, New-York, pp. 387-393. Ashwal, L.D., Morgan, P., Kelley, S.A., Percival, J., 1987. Heat production in an Archean crustal profile and implications for heat flow and mobilization of heat producing elements. Earth Planet. Sci. Lett. 85, 439-450. Ballard, S., Pollack, H.N., Skinner, N.J., 1987. Terrestrial heat flow in Botswana and Namibia. J. Geophys. Res. 92, 62916300. Beck, A.E., Dharba, D.M., Schloessin, H.H., 1978. Lattice conductivities of single crystal and polycrystalline materials at mantle pressures and temperatures. Phys. Earth Planet. Int. 17, 35-53. Birch, F., Roy, E.R., Decker, E.R., 1968. Heat flow and thermal history in New England and New York. In: An-Zen, E. (Ed.), Studies of Appalachian Geology. Wiley (Interscience), New York, pp. 437-451. Christensen, N.I., Mooney, W.D., 1995. Seismic velocity structure and composition of the continental crust: a global view. J. Geophys. Res. 100, 9761-9788. Clauser, C., Gieses, P., Huenges, E., Kohl, T., Lehmann, H., Rybach, L., Safanda, J., Wilhelm, H., Windlow, K., Zoth, G., 1997. The thermal regime of the crystalline continental crust: implications from the KTB. J. Geophys. Res. 102, 1841718441. Clowes, R.M., Cook, F.A., Green, A.G., Keen, C.E., Ludden, J.N., Percival, J.A., Quinlan, G.M., West, G.F., 1992. LITHOPROBE New perspectives on crustal evolution. Can. J. Earth Sci. 29, 1831-1864. Cull, J.P., 1991. Heat flow and regional geophysics in Australia. In: Cermak, V., Rybach, L. (Eds.), Terrestrial Heat Flow and the Lithosphere Structure. Springer Verlag, Berlin, pp. 486500. Davaille, A., Jaupart, C., 1994. Onset of thermal convection in fluids with temperature dependent viscosity: application to the oceanic mantle. J. Geophys. Res. 99, 19853-19866. Doin, M.P., Fleitout, L., Christensen, U.R., 1997. Mantle convection and stability of depleted and undepleted continental lithosphere. J. Geophys. Res. 102, 2771-2788. Drury, M.J., 1985. Heat flow and heat generation in the Churchill Province of the Canadian Shield and their paleotectonic significance. Tectonophysics 115, 25-44. Drury, M.J., Taylor, A.E., 1987. Some new measurements of heat flow in the Superior Province of the Canadian Shield. Can. J. Earth Sci. 24, 1486-1489. Drury, M.J., Jessop, A.M., Lewis, T.J., 1987. The thermal nature of the Canadian Appalachians. Tectonophysics 113, 1-14. Durrheim, R.V., Green, R.W.E., 1992. A seismic refraction investigation of the Archean Kaapvaal craton, South Africa, using mine tremors as the energy source. Geophys. J. Int. 108, 812-832.
1 0T
02T
A
Ot
0Z 2
K exp'
K
(B1)
at),
where K is the thermal diffusivity, A is the heat source density, K is the thermal conductivity, and ol ~ 0 . 5 1 0 . 9 year -1. We assume that the temperature is 0 at the surface z - 0 and that the heat flow is 0 at the base of the lithosphere z - L. We shall assume that the effect of the initial conditions has been damped out and can be neglected, and a solution can be obtained by expanding the temperature in the following Fourier series:
T ( z , t ) = }7~ c~(t)sin (2k + 1 ) ~
(B2)
k=0
with 1 ~ s i n 1 = - 27 r k = 0 2 k + l
2k+l)
-~
0
The solution the temperature is obtained as: 16AL 2
r(z,t)
-
. •
"n-pC e x. p ( a t.)
. exp( 2-
k=0
(2k +
1)2 ,
2 t/4L2)
+ 1)2=2 (B4)
Neglecting the transient terms due to the poorly defined initial condition, we obtain the surface heat flow as: OT ~ 1 K - - = 8 AL exp( - c~t) E 0z ~=0 4 a t - (2k -+- 1) 277"2
(BS) aT tan~/ar K-- =AL exp( - a t ) Oz
V/c~r
with r = L2/K.
(B6)
C. Jaupart, J.C. Mareschal / Lithos 48 (1999) 93-114
Eade, K.E., Fahrig, W.F., 1971. Geochemical evolutionary trends of continental plates. A preliminary study of the Canadian Shield. Geol. Surv. Canada Bull. 179, 59. Fountain, D.M., 1986. Is there a relationship between seismic velocity and heat production for crustal rocks?. Earth Planet. Sci. Lett. 79, 145-150. Fountain, D.M., Salisbury, M.H., Furlong, K.P., 1987. Heat production and thermal conductivity of rocks from the Pikwitonei-Sachigo continental cross section, central Manitoba: implications for the thermal structure of Archean crust. Can. J. Earth Sci. 24, 1583-1594. Gao, S., Luo, T.-C., Zhang, B.-R., Zhang, H.-F., Han, Y.-W., Hu, Y.-K., Zhao, Z.-D., 1998a. Chemical composition of the continental crust as revealed by studies in East China. Geochim. Cosmochim. Acta 62, 1959-1975. Gao, S., Zhang, B.-R., Jin, Z.-M., Kern, H., Luo, T.-C., Zhao, Z.-D., 1998b. How mafic is the lower continental crust?. Earth Planet. Sci. Lett. 161, 101-117. Green, R.W.E., Durrheim, R.J., 1990. A seismic refraction investigation of the Namaqualand metamorphic complex, South Africa. J. Geophys. Res. 95, 19927-19932. Guillou, L., Mareschal, J.-C., Jaupart, C., Gari@y, C., Bienfait, G., Lapointe, R., 1994. Heat flow gravity and structure of the Abitibi belt, Superior Province, Canada: implications for mantle heat flow. Earth Planet. Sci. Lett. 122, 103-123. Guillou-Frottier, L., Mareschal, J.-C., Jaupart, C., Gari@y, C., Lapointe, R., Bienfait, G., 1995. Heat flow variations in the Grenville Province, Canada. Earth Planet. Sci. Lett. 136, 447-460. Guillou-Frottier, L., Jaupart, C., Mareschal, J.-C., Gari@y, C., Bienfait, G., Cheng, L.Z., Lapointe, R., 1996. High heat flow in the Thompson Belt of the Trans-Hudson Orogen, Canadian Shield. Geophys. Res. Lett. 23, 3027-3030. Gupta, M.L., Sundar, A., Sharma, S.R., 1991. Heat flow and heat generation in the Archean Dharwar cratons and implications for the Southern Indian Shield geotherm and lithospheric thickness. Tectonophysics 194, 107-122. Gupta, M.L., Sundar, A., Sharma, S.R., Singh, S.B., 1993. Heat flow in the Bastar craton, central Indian Shield: implications for thermal characteristics of Proterozoic cratons. Phys. Earth Planet. Int. 78, 23-31. Gurnis, M., 1988. Large-scale mantle convection and the aggregation and dispersal of supercontinents. Nature 332, 695-699. Hart, R.J., Nicolaysen, L.O., Gale, N.H., 1981. Radioelement concentrations in deep profile through Archean basement of the Vredefort structure. J. Geophys. Res. 86, 10639-10652. Hart, S.R., Steinhart, J.S., Smith, T.J., 1994. Terrestrial heat flow in Lake Superior. Can. J. Earth Sci. 31,698-708. Hoffman, P.F., 1989. Precambrian geology and tectonic history of North America. In: Bally, A.W., Palmer, E.R. (Eds.), The Geology of North America--An Overview. GSA, Boulder, CO, pp. 447-512. Jaupart, C., 1983. Horizontal heat transfer due to radioactivity contrasts: causes and consequences of the linear heat flow-heat production relationship. Geophys. J. R. Astron. Soc. 75, 411435. Jaupart, C., Mann, J.R., Simmons, G., 1982. A detailed study of
113
the distribution of heat flow and radioactivity in New Hampshire (USA). Earth Planet. Sci. Lett. 59, 267-287. Jaupart, C., Mareschal, J.-C., Guillou-Frottier, L., Davaille, A., 1998. Heat flow and thickness of the lithosphere in the Canadian Shield. J. Geophys. Res. 103, 15269-15286. Jessop, A.M., 1971. The distribution of glacial perturbation of heat flow in Canada. Can. J. Earth Sci. 8, 162-166. Jessop, A.M., Lewis, T.J., Judge, A.S., Taylor, A.E., Drury, M.J., 1984. Terrestrial heat flow in Canada. Tectonophysics 103, 239-261. Jones, M.Q.W., 1987. Heat flow and heat production in the Namaqua mobile belt, South Africa. J. Geophys. Res. 92, 6273-6289. Jones, M.Q.W., 1988. Heat flow in the Witwatersrand Basin and environs and its significance for the South African shield geotherm and lithosphere thickness. J. Geophys. Res. 93, 3243-3260. Jones, M.Q.W., 1992. Heat flow anomaly in Lesotho: implications for the southern boundary of the Kaapvaal craton. Geophys. Res. Lett. 19, 2031-2034. Jordan, T.H., 1981. Continents as a chemical boundary layer. Trans. R. Soc. (London), Ser. A 301,359-373. Lachenbruch, A.H., 1970. Crustal temperature and heat production: implications of the linear heat flow heat production relationship. J. Geophys. Res. 73, 3292-3300. Lenardic, A., Kaula, W.M., 1996. Near surface thermal/chemical boundary layer convection at infinite Prandtl number: two-dimensional numerical experiments. Geophys. J. Int. 126, 689711. Mareschal, J.-C., Pinet, C., Gari@y, C., Jaupart, C., Bienfait, G., Dalla-Coletta, G., Jolivet, J., Lapointe, R., 1989. New heat flow density and radiogenic heat production data in the Canadian Shield and the Qu6bec Appalachians. Can. J. Earth Sci. 26, 845-852. Mareschal, J.-C., Jaupart, C., Cheng, L.Z., Rolandone, F., Gariepy, C., Bienfait, G., Guillou-Frottier, L., Lapointe, R., 1999. Heat flow in the Trans-Hudson Orogen of the Canadian Shield: implications for Proterozoic continental growth, J. Geophys. Res., in press. McKenzie, D.P., Bickle, M.J., 1988. Volume and composition of melt generated by extension of the lithosphere. J. Petrol. 29, 625-679. Moresi, L.-N., Lenardic, A., 1997. Three-dimensional numerical simulations of crustal deformation and subcontinental mantle convection. Earth Planet. Sci. Lett. 150, 233-243. Morgan, P., 1985. Crustal radiogenic heat production and the selective survival of continental crust. J. Geophys. Res. 90, C561-C570. Nicolas, A., J.-P. Poirier, 1976. Crystalline Plasticity and Solid State Flow in Metamorphic Rocks. Wiley, New York. Nicolaysen, L.o., Hart, R.J., Gale, N.H., 1981. The Vredefort Radioelement Profile extended to supracrustal strata at Carletonville, with implications for continental heat flow. J. Geophys. Res. 86, 10653-10661. Nixon, P.H., Rogers, N.W., Gibson, I.L., Grey, A., 1981. Depleted and fertile mantle xenoliths from southern African kimberlites. Annu. Rev. Earth Planet. Sci. 9, 285-309.
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Nyblade, A.A., Pollack, H.N., 1993. A global analysis of heat flow from Precambrian terrains: implications for the thermal structure of Archean and Proterozoic lithosphere. J. Geophys. Res. 98, 12207-12218. Parsons, B., McKenzie, D.P., 1978. Mantle convection and the thermal structure of the plates. J. Geophys. Res. 83, 44854495. Pinet, C., Jaupart, C., 1987. The vertical distribution of radiogenic heat production in the Precambrian crust of Norway and Sweden: geothermal implications. Geophys. Res. Lett. 14, 260-263. Pinet, C., Jaupart, C., Mareschal, J.-C., Gari6py, C., Bienfait, G., Lapointe, R., 1991. Heat flow and structure of the lithosphere in the eastern canadian shield. J. Geophys. Res. 96, 1994119963. Polyak, B.G., Smirnov, Y.B., 1968. Relationship between terrestrial heat flow and tectonics of the continents. Geotectonics 4, 205-213. Qiu, X., Priestley, K., McKenzie, D., 1996. Average lithospheric structure of southern Africa. Geophys. J. Int. 127, 563-587. Rudnick, R.L., Fountain, D.M., 1995. Nature and composition of the continental crust: a lower crustal perspective. Rev. Geophys. 33, 267-309. Rudnick, R.L., McDonough, W.F., O'Connell, R.J., 1998. Thermal structure, thickness and composition of continental lithosphere. Chemical Geology 145, 395-411. Sass, J.H., A.H. Lachenbruch, 1979. Thermal regime of the Australian continental crust. In: McElhinny, M.W. (Ed.), The
Earth: its Origin, Structure and Evolution. Academic Press, New York, pp. 301-351. Sch~irmeli, G., 1979. Identification of radiative thermal conductivity in olivine up to 25 kbar and 1500 K, Proc. 6th AIRAPT Conf. In: Timmerhauf, K.D., Barber, M.S. (Eds.). Plenum, New York, pp. 60-74. Schatz, J.F., Simmons, G., 1972. Thermal conductivity of earth materials. J. Geophys. Res. 77, 6966-6983. Sclater, J.G., Parsons, B., Jaupart, C., 1981. Oceans and continents: similarities and differences in the mechanisms of heat loss. J. Geophys. Res. 86, 11535-11552. Shaw, D.M., Cramer, J.J., Higgins, M.D., Truscott, M.G., 1986. Composition of the Canadian Precambrian Shield and the continental crust of the earth. In: Dawson, J.B. et al. (Eds.), Nature of the Lower Continental Crust. Geol. Soc., London, pp. 257-282. Smithson, S.B., Decker, E.R., 1974. A continental crustal model and its geothermal implications. Earth Planet. Sci. Lett. 22, 215-225. Smithson, S.B., Ramberg, I.B., 1979. Gravity interpretation of the Egersund anorthosite complex, Norway: its petrological and geothermal significance. Bull. Geol. Soc. Am. 90, 199-204. Taylor, S.R., McLennan, S.M., 1985. The continental crust: its composition and evolution. Blackwell, Cambridge, MA, 312 pp. Vitorello, I., Pollack, H.N., 1980. On the variation of continental heat flow with age and the thermal evolution of continents. J. Geophys. Res. 85, 983-995.
LITHOS
ELSEVIER
Lithos 48 (1999) 115-133
Stability and dynamics of the continental tectosphere Steven S. Shapiro ~, Bradford H. Hager * , Thomas H. Jordan Department of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA 02139, USA Received 19 October 1998; received in revised form 8 February 1999; accepted 15 February 1999
Abstract Continental cratons overlie thick, high-viscosity, thermal and chemical boundary layers, where the chemical boundary layers are less dense than they would be due to thermal effects alone, perhaps because they are depleted in basaltic constituents. If the continental tectosphere is the same age as the overlying Archaean crust, then the continental tectosphere must be able to survive for several billion years without undergoing a convective instability, despite being both cold and thick. Since platforms and shields correlate only weakly with Earth's gravity and geoid anomalies, acceptable models of the continental tectosphere must also satisfy this gravity constraint. We investigate the long-term stability of the continental tectosphere by carrying out a number of numerical convection experiments within a two-dimensional Cartesian domain. We initiate our experiments with a tectosphere (thermal and chemical boundary layers) immersed in a region of uniform composition, temperature, and viscosity, and consider the effects on the stability of the tectosphere of (1) activation energy (used to define the temperature dependence of viscosity), (2) compositional buoyancy, and (3) linear or non-linear rheology. The large lateral thermal gradients required to match oceanic and tectosphere structures initiate the dominant instability, a "drip" which develops at the side of the tectosphere and moves to beneath its center. High activation energies and high background viscosities restrict the amount and rate of entrainment. Compositional buoyancy does not significantly change the flow pattern. Rather, compositional buoyancy slows the destruction process somewhat and reduces the stress within the tectosphere. With a non-Newtonian rheology, this reduction in stress helps to stiffen the tectosphere. In these experiments, dynamical systems that adequately model the present ocean-continent structures have activation energy E* > 180 kJ m o l e a value about one third the estimate of activation energy for olivine, E* -- 520 kJ mole-1. Although for E* ~ 520 kJ mole-1, compositional buoyancy is not required for the tectosphere to survive, the joint application of longevity and gravity constraints allows us to reject all models not containing compositional buoyancy, and to predict that the ratio of compositional to thermal buoyancy within the continental tectosphere is approximately unity. 9 1999 Published by Elsevier Science B.V. All rights reserved. Keywords: Continental tectosphere; Dynamics; Conductive cooling; Mantle convection; Mantle rheology
* Corresponding author i Now at the Department of Physics, Guilford College, Greensboro, NC 27410, USA. 0024-4937/99/$- see front matter 9 1999 Published by Elsevier Science B.V. All rights reserved. PII: S 0 0 2 4 - 4 9 3 7 ( 9 9 ) 0 0 0 2 5 - 0
116
S.S. Shapiro et al./Lithos 48 (1999) 115-133
1. Introduction The weak association between platforms and shields and long-wavelength geoid height anomalies (Shapiro et al., 1999), as well as the near constancy of continental freeboard (e.g., Wise, 1974) suggest that the continental tectosphere did not form simply by conductive cooling. Jordan (1975, 1978) proposed that the thick, cold, continental chemical boundary layer (CBL) inferred from the analysis of seismic data (see also, for example, more recent studies by Grand, 1994; Su et al., 1994; EkstriSm and Dziewonski, 1995; Masters et al., 1996) is stabilized against convective disruption by compositional variations that yield neutral buoyancy (i.e., the continental tectosphere and neighboring oceanic material have the same density profile). These compositional variations have been attributed to a depletion, relative to that from the source of mid-ocean-ridge volcanism, of the denser basaltic constituents (garnet and clinopyroxene) in the continental tectosphere (Jordan, 1978). Monte Carlo simulations indicate that the strong association between platforms and shields and upper-mantle shear-wave anomalies is not simply fortuitous (Shapiro, 1995). This global relationship supports the hypothesis (Jordan, 1975) that these thick CBLs translate coherently with continental plate motions. Combining this conjecture with the measured ages of South African diamond inclusions (e.g., Richardson et al., 1984) and ages obtained from r h e n i u m - o s m i u m and other isotope systematics
(Walker et al., 1989; Pearson et al., 1995) supports the supposition that the continental tectosphere could remain intact in the convecting mantle for times in excess of a billion years (Jordan, 1978). If the tectosphere is to survive in a convecting mantle, it must be stable both to double diffusive instabilities resulting from compositional buoyancy (e.g., Stevenson, 1979) and to tractions from the convecting mantle in which it is immersed. In this study, we investigate primarily the former. Our aim is to quantify the role of viscosity and compositional buoyancy in determining whether a tectosphere would undergo spontaneous disintegration. The boundary conditions in our model are chosen to suppress thermal anomalies associated with the background large-scale mantle convection. We consider the effects of activation energy, compositional buoyancy, and dependence of rheology on stress. We further constrain the range of acceptable parameters by requiting that they not only produce long-term CBL stability, but also create density distributions that yield geoid height anomalies consistent with those observed over platforms and shields (Shapiro et al., 1999).
2. Numerical formulation We use a double-diffusive version of the finiteelement program, ConMan (King et al., 1990), to
600
(km)
(%1
x (km)
Fig. 1. Contours of initial composition (left) and temperature (mirrored -- fight). Dimensionless composition is contoured in increments of 0.1, with C = 1 at the top and C = 0.1 at the base of the CBL. Dimensionless temperature is contoured in increments of 0.1, with T = 0 at the top and T = 1 at the base of the TBL. The C = T = 0.1 contours are thick. Fluctuations in the initial temperature field are due to the superposition on the temperature field of a white noise perturbation with a zero mean and a peak-to-peak amplitude of 0.01. Composition is determined from the unperturbed temperature field according to Eq. (9). The temperature at the base of the CBL, TCBL,is 1170~ {0.9}. The center frame displays, as a function of depth, the difference between the average mid-tectosphere (x < 200 km) density and the average lateral density (see Eq. (10)) for B = 0 (dashed line), B = 1 (thin line), and B = 1.5 (thick line).
117
S.S. Shapiro et al./Lithos 48 (1999) 115-133
solve numerically the advection-diffusion equations for flow of an incompressible, infinite Prandtl number fluid in a two-dimensional Cartesian domain. With two fields affecting density" temperature, T, and composition, C, the relevant (dimensionless) equations are those of momentum balance: =
Vp
-
Ra (;r + 8C)
(1)
continuity: Vu =0
(2)
where E* is the activation energy, Toff is the offset for dimensionless surface temperature required for dimensionless T = 0 to correspond to the = 273 K surface temperature of Earth. The background viscosity, ~7~, implicitly accounts for pressure variations, including the effects of phase changes. A generic form for non-Newtonian viscosity is r/NN = ~Tr~(~.0/r)n- 1, where ~- is the second invariant of the stress tensor, TO is the stress value when the Newtonian and non-Newtonian viscosities are equal, and n, the power law exponent, is assumed to be three for
conservation of energy (with no internal heating): OT
= - u V ; r + v2;r
(3) ]
and its compositional analog"
/
I I
0C 0t
1 = -uVC+
Le
(4)
VZc
where r/ is dynamic viscosity, u velocity, p pressure, t time, Le Lewis number (ratio of thermal to compositional diffusivity), and ~ the unit vector in the direction of increasing depth. The thermal Rayleigh number is Ra T - c ~ g p A T d 3 / K T ~ ) l where ce is the coefficient of thermal expansion, g the acceleration due to gravity, p the density, AT the difference between the temperature at the bottom and that at the top of the domain of depth d, KT the thermal diffusivity, and r/~ a reference value of dynamic viscosity we define to be the viscosity corresponding to dimensionless T = 1. The buoyancy number, B =- A p c / p c e A T , with A Pc the difference in density due to a unit change in composition, describes the ratio of compositional to thermal buoyancy. Unlike thermal gradients, which evolve through the diffusion of heat over geologic time scales, compositional gradients are essentially unaffected by solid-state diffusion; hence Le is effectively infinite. Owing to numerical constraints, however, we are limited to Le < 100 (Brooks, 1981). We define a temperature-dependent Newtonian viscosity, r/N, using an Arrhenius law (e.g., King, 1990): E* r/N ( T ) = ~71 exp T + Toff
-- f-------I "Jr-To7
(5)
i
|
"''I""
~
I
500
,
1020
,
"i;25 viscosity (Pa s)
Fig. 2. Newtonian viscosity profiles based on background viscosities (a) NLO (thick line) and (b) HGPA (thick line), and evaluated using the initial temperature field of the mid-tectosphere and three activation energies: E*f (dash-dotted lines), E*f/3 (dashed lines), and E*f/9 (thin lines), where E*f = 522 kJ mole-1. The dotted lines represent the viscosity profile corresponding to the initial mid-oceanic temperature profile. Note: we do not model NLO's 100-fold increase in viscosity at 670 km depth because its effect on our experiments is insignificant (Shapiro, 1995).
118
s.s. Shapiro et al./Lithos 48 (1999) 115-133
our n o n - N e w t o n i a n ex p e r i m e n t s . Since strain rates are additive, an effective viscosity, ~eff, that in the limit of low stress yields N e w t o n i a n creep and in the limit of high stress yields n o n - N e w t o n i a n creep, can be written as r/eff = [r/~ 1 + r/~ 1 ] - 1 . W e use a value of ~'0 = 1 M P a (10 bars) to define the transition b e t w e e n N e w t o n i a n and n o n - N e w t o n i a n r h e o l o g y and, to avoid n u m e r i c a l p r o b l e m s , we fix the maxim u m d i m e n s i o n l e s s viscosity at ~ c u t / ~ l , w h e r e '0cut = 10 5. B e c a u s e this viscosity cutoff reduces the strength of the u p p e r m o s t mantle, we simulate the effect of this strong layer by i m p o s i n g no-slip b o u n d a r y conditions at the surface.
Table 2 Non-Newtonian rheology: experimental parameters and resulting (i) geoid height anomalies (gN) over the tectosphere, and (ii) depth of the base of the tectosphere CBL (ZcBL)- Both quantities are calculated at 1000 My. We keep all entries in each column to the same place for ease of reading even though not all digits are significant. We do not list the results with E* = 80 kJ molebecause these calculations result in such a rapid destruction of the tectosphere that assigning a thickness is meaningless. See Fig. 8 for a graphical representation of this table E*
B
"Ob (z)
g N (m)
ZCBL (km)
0.0
HGPA NLO HGPA NLO HGPA NLO HGPA NLO HGPA NLO HGPA NLO
- 10 _+5 - 17 _+11 - 12 _+7 - 2 _+5 - 2_ 1 10 _ 6 - 68 +_42 - 83 _ 48 - 11 _+9 - 19 + 11 21 _+10 14 +_8
92 138 304 320 289 271 285 320 329 348 340 348
(kJ mole- 5)
174
1.0 1.5
3. Experimental design and parameters
522
In this study, w e address the stability of a C B L e x p o s e d to c o n v e c t i v e and b u o y a n c y stresses while not explicitly including all of the effects of a convecting mantle. M o d e l i n g n u m e r i c a l l y a d o m a i n the size of Earth with a realistic R a y l e i g h n u m b e r , and t e m p e r a t u r e and s t r e s s - d e p e n d e n t viscosity, is computationally intensive e v e n in t w o - d i m e n s i o n s be-
Table 1 Newtonian rheology: experimental parameters and resulting (i) geoid height anomalies (gN) over the tectosphere, and (ii) depth of the base of the tectosphere CBL (ZcBL). Both quantities are calculated at 1000 My. We keep all entries in each column to the same place for ease of reading even though not all digits are significant. We do not list the results with E* = 80 kJ mole-1 because these calculations result in such a rapid destruction of the tectosphere that assigning a thickness is meaningless. See Fig. 8 for a graphical representation of this table E* (kJ mole- 1)
B
"i"]b (z)
g N (m)
ZCBL (km)
174
0.0
HGPA NLO HGPA NLO HGPA NLO HGPA NLO HGPA NLO HGPA NLO
- 37 _+28 - 62 _+40 5 +_3 - 8 _+5 14 +_8 21 _+10 - 131 +_86 - 119 _+70 - 53 + 34 - 4 2 _+20 - 13 _+8 - 3 _+5
216 250 284 287 297 302 317 309 346 346 346 346
1.0 1.5 522
0.0 1.0 1.5
0.0 1.0 1.5
cause of the high spatial and t e m p o r a l resolution required to m o d e l accurately the flow globally. Our t ect osphere is i m m e r s e d in a hot (and h e n c e r e d u c e d viscosity) i sot hermal and i s o c h e m i c a l e n v i r o n m e n t . W e obtain basal tractions in the fraction of a m e g a pascal range, c o m p a r a b l e to estimates for intraplate sublithospheric tractions on Earth (e.g., H a g e r and O ' C o n n e l l , 1981), so this a p p r o x i m a t i o n s e e m s reasonable. (Stresses in subduction zone e n v i r o n m e n t s are likely s o m e w h a t higher, but, as di scussed b e l o w , regions with p r e s e r v e d t ect osphere typically h a v e b e e n far f r o m subduction zones for m o s t of their history.) W e solve our s y s t e m of equations in a t w o - d i m e n sional Cartesian d o m a i n using a 76 • 38 grid of square e l e m e n t s (see Shapiro (1995) for a discussion of the sensitivity of the results to grid resolution, aspect ratio, and symmet ry). W e begin all of our e x p e r i m e n t s with an " o c e a n i c " t hermal b o u n d a r y layer ( T B L ) of 100 k m {0.13} and a " c o n t i n e n t a l " T B L of 400 k m thickness {0.53} a b o v e a region of u n i f o r m t e m p e r a t u r e (Fig. 1). (To facilitate rescaling our m o d e l results to other plausible initial conditions with thinner or thicker T B L ' s, we present d i m e n s i o n -
S.S. Shapiro et al./Lithos 48 (1999) 115-133
t
0 My; Vmax
34 cm/y,r
119
tOtect-11" Obase--2 (mW/m2)
9
,'
/
l
l
600 t
l
t
l
I
'
150 My; V m a x - 7 cn~,yr
0
/
.Qtect
t = 500 My; V,m a x = 9 cm/,yr
o
~
~
.,.
13;
Qbase=
10 (m,W/m2)
Qtect = 18" Qbase = 5
.
.
.
(m~//m
(C)
.
%
600 |
I
t = 1000 My; Vmax = 10 cm/yr _
f
,Q)tect = 20; Obase = 9 (m~y/m 2)
(d)
Y
600
x (kin)
(%)
Fig. 3. Parameters: E*f = E*f/3, B = 0, n = 3, r/b(Z) = NLO. Four frames: t = (a) 0, (b) 150, (c) 500, and (d) 1000 My. Each frame contains (left) contours of composition (purely a tracer field having no effect on the dynamics of the fluid) with superposed velocity arrows, (center) the difference, as a function of depth, between the average mid-tectosphere (x < 200 km) density and the average lateral density (see Eq. (10)), and (mirrored right) contours of temperature with superposed velocity arrows. Contour levels as in Fig. 1.
120
S.S. Shapiro et al./Lithos 48 (1999) 115-133
less parameter values within brackets.) We apply the following boundary conditions along the top (z = 0): ux=0
(6a)
uz=O
(6b)
T = 0~ {0}
(6c)
OC
-0 Oz along the bottom (z = 760 km {1.0}),
(6d)
rzz=O ux=O
(7a) (7b)
0T --
=0
Oz OC
(7c)
-0 Oz along the sides ( x - 0, 1520 km {2.0}),
(7d)
~-xz-O
(8a)
ux=0
(8b)
OT Ox OC
=o
senting an ocean and the other a continental tectosphere. We obtain our initial temperature field by solving the heat conduction equation subjected to the above temperature boundary conditions and a fixed temperature, T = 1300~ {1.0}, along the base ( z 400 km {0.53}) and side (x < 800 km {1.05}) of the tectosphere TBL. We remove these supplementary boundary conditions once the initial temperature field is formed. The compositional field represents the degree of basalt depletion with respect to the oceanic average, with a higher value of C representing a larger amount of depletion and yielding a lower normative density. Following the hypothesis of Jordan (1978, 1988) that, within the continental tectosphere, contours of constant composition are parallel to isotherms, we create an initial (dimensionless) composition field (Fig. 1) that is a function of the initial (dimensionless) temperature field:
C=I-T,
0~
C = 0,
T > TCB L
T~
TCB L
(9)
(8c)
=0 (8d) Ox and beneath the oceanic TBL (x > 800 km {1.05}; z = 100 km {0.13}): The thermal boundary condition at the bottom of the domain and at the base of the oceanic lithosphere were designed to suppress hot plumes and small-scale convection beneath oceanic regions, respectively. Our initial temperature field (Fig. 1) contains two adjoining TBLs of different thicknesses, one repre-
where TcBL = 1170~ {0.9} defines the temperature at the base of the CBL. It is not essential that TcBL have this particular value. Somewhat larger values lead to initial conductive thickening of the TBL while somewhat smaller values lead to the lower part of the TBL dropping off into the fluid below (see Shapiro (1995)). We investigate buoyancy ratios of 0, 1, and 1.5, where B - 0 corresponds to density unaffected by composition, B - 1 to the isopycnic hypothesis of Jordan (1988), and B - 1.5 to a tectosphere with net positive buoyancy.
Fig. 4. Model parameters as in Fig. 3. (a) Geoid height anomalies (gN) at t = 0 (dash-dotted line), 100 (thin solid line), and 1000 (thick solid line) My. (b) Dynamic topography (h) at t-- 0 (dash-dotted line), 100 (thin solid line), and 1000 (thick solid line) My. (c) Viscosity field (7) with superposed velocity arrows (left) and second invariant of the stress tensor (~'(II)) (mirrored fight) at t = 0. Viscosity contours are spaced by factors of 100, with the thick line representing the lowest contour level (102~ Pa s). For the stress field, the thick line represents the lowest contour level (0.5 MPa) and each succeeding contour indicates a stress value a factor of two larger than that for the immediately preceding contour. (d) Viscosity field (7) with superposed velocity arrows (left) and second invariant of the stress tensor (~'(II)) (mirrored fight) at t = 1000 My. Contour intervals as in (c). (e) Initial composition field subtracted from the composition field at t-- 1000 My (left) and initial temperature field subtracted from the temperature field at t-- 1000 My (mirrored fight). Dimensionless contours are spaced in increments of 0.1 with dashed lines representing a loss of composition/temperature and thick solid lines representing a gain. The zero contours are shown with thin solid lines. (f) Area (A) of the - 0 . 1 difference contour, normalized by the area of the initial oceanic and tectosphere CBL, representing a loss of composition (thin line, open circles), and the change (A ZCBL) in the depth of the base of the tectosphere CBL (C-- 0.1) at 1000 My expressed as a percentage of the initial depth (thick line, asterisks). The depth of the base of the tectosphere CBL is estimated from the median depth of the C = 0.1 contour within the tectosphere CBL. (g) Conductive heat flux (Q) through the surface of the tectosphere (thin line, open circles) and advective heat flux through the base of the domain (thick line, asterisks).
(,q~) ~
(,~)
S
. zv :V
(O)a~ z / ~
x
(t) x
00~,'~ 00~ (o)•
ooot)J/
',o)~ - (,{~ d o o t ) 5
' ;
t
~ '~, ,, \ . .
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/
'
l /
l
,
,
"~
,
00g
(IDa,
EL
~{/mb OI
9
,
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|
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t
,
'
I
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,
r
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00~,'~ 00g 0
S
m
oB .
.
(q)
[~I
.
.
.
.
.
.
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.
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.
.
.
.
t
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/ 7~ la o.qdmt5; "~'~
s.s. Shapiroet al./Lithos 48 (1999) 115-133
122
For background viscosity profiles T~b(Z), we use the models H G P A (Hager, 1991) and NLO (Nakada and Lambeck, 1989) (Fig. 2). We base our selection of activation energies on the estimate for dry olivine of E* = E~*ef- 522 kJ mole-1 by Ashby and Verrall (1977); we use E* =E~*f, E*f/3, and E*f/9 to explore the sensitivity of tectosphere stability to activation energy (Tables 1 and 2). Fig. 2 shows, for each activation energy used, the corresponding initial mid-oceanic and mid-tectosphere Newtonian viscosity profiles. In all of our numerical experiments, we use representative values for the following quantities" ce = 3 • 1 0 - 5 ~ l, g = 9.8 m s-2, p = 3.5 x 10 3 kg m -3 and KT = 10 -6 m 2 s -1
4. Buoyancy contrast One can begin to analyze the stability of a particular thermal and compositional structure by calculating the continent-ocean buoyancy contrast, i.e., by calculating at each depth the difference in density between the average mid-tectosphere ( x _< 200 km) value and the laterally averaged value (Fig. 1):
=- ceAT[(T+ BC>lx<_2OOkm--
BC>x]
P (10) For example, a negative value of ~ p ( z ) / p indicates that at depth z, the central region of the tectosphere is lighter, on average, than the average of the ocean and continent values. Considering ~ p / p as a function of depth allows one to predict whether a structure might tend to remain near its initial configuration. As one can see from the plots of ~ p ( z ) / p for B = 0 and B = 1 in Fig. 1, a structure with B = 1 has a greater chance for survival than a structure with B = 0 because, with B = 0, the mid-tectosphere region is much denser than the surrounding material. Of course, this plot neither tells us what structures, if
any, are stable nor how an unstable structure might disintegrate. In the subsequent discussion, we take 1000 My as the characteristic time at which to assess stability. By this time the CBL has either been destroyed or the fluid flow is sufficiently regular that one can reliably make predictions concerning longterm stability.
5. Compositional buoyancy: effect on boundary layer stability To illustrate the effect of composition-induced buoyancy on boundary-layer stability, we discuss in detail the evolution of two cases that differ only in the value of B. Since an acceptable model of the continental tectosphere must satisfy both the longevity and the gravity constraints, at 1000 My we calculate the depth to the base of the CBL, ZCBL, and estimate the associated geoid height anomaly, 8 N (Tables 1 and 2). In the first example, we take E'f~3, B = 0, and n = 3, which leads to very rapid destruction of the tectosphere CBL. With B = 0, composition is simply a tracer field it has no effect on the motion of the fluid. At 1000 My, the CBL is essentially gone, having been washed away by the flow driven by the lateral variations in density (Fig. 3). The tectosphere-ocean buoyancy contrast, ~ p ( z ) / p , decreases in magnitude and, as the tectosphere CBL disappears, becomes non-zero only at shallow depths. The average geoid height anomaly associated with the beginning of the experiment is much greater in amplitude than those observed over platforms and shields (e.g., Shapiro et al. (1999)) (Fig. 4a). Of course, as the source of the density contrast, the tectosphere TBL disappears, the average geoid height anomaly decreases accordingly. Similarly, the variation in dynamic topography is unreasonably large when compared with the near constancy through
Fig. 5. (a) Parameters: E* = E'f/3, B = 0, n = 3, Tlb(Z) = NLO (Fig. 3). Viscosity field ('q) with superposed velocity arrows (left) and second invariant of the stress tensor (~-(II)) (mirrored right)at t = 0. Velocity scaling is clipped at Ucut = 1 cm/year to exhibit lower values more clearly. Contour intervals as in Fig. 4c. (b) Viscosity field (r/) with superposed velocity arrows (left) and second invariant of the stress tensor (-r(II)) (mirrored right) at t= 1000 My. Velocity scaling as in (a). Contour intervals as in Fig. 4c. (c) Parameters: E* = E*f/3, B-- 1.0, n = 3, ~Tb(Z)= NLO (Fig. 6). Viscosity field (r/) with superposed velocity arrows (left) and second invariant of the stress tensor (~'(II)) (mirrored right) at t = 0. Velocity scaling as in (a). Contour intervals as in Fig. 4c. (b) Viscosity field (r/) with superposed velocity arrows (left) and second invariant of the stress tensor (-r(II)) (mirrored - - right) at t = 1000 My. Velocity scaling as in (a). Contour intervals as in Fig. 4c.
123
S.S. Shapiro et al./Lithos 48 (1999) 115-133
the tectosphere much of the flow occurs in the Newtonian regime, although a substantial amount also occurs in regions of high stress (Fig. 5a and b). At 1000 My, the tectosphere is gone, so the size of
time of Earth's continental freeboard (e.g., Wise, 1974) (Fig. 4b). The initial viscosity and stress fields (Fig. 4c) show that the transition between Newtonian and non-Newtonian rheology occurs near the base of
0 I; O; Vclip -..--
...~
-c(II)
1 cm/yr
,,,.~ .,,,.~ .,.,,~
~
"
,
,
,, ~, , , . . . . . i . _ ~ , , ~ ~ . . - ~ ~ ~
,~t171/,,"
xii)~x~t
~ = 1000 My;,Vclip = 1 cm/yr 20O .
,
t
I
t
t...,'...,,...,...,
,
B
1}=O; Vclip = 1 cm/yr
E I tilll~~,,,.....
I
" ""//r ''
0
~
'"t
I
..\\\ \ ~ \\\~ t 1
l,J I ,
~ - - ~ - ~
< < <, i
600
"~ "" ,: ,~ < i K K i ~/~'s "~C"i /"~'-_
x (km)
x (kin)
124
S.S. Shapiro et al./Lithos 48 (1999) 115-133
"Vmax=~m/yr
.
"
150 My; V m a x
4 crn(,yr
.Qtect
3 (m~y/m
11; Qbase
N
600
1
t = 500 My; V m a x = 6 cm(,yr
tect-" 12"
base = ~ W l m .
,
0,'
t = 1000 My; V m a x = 3 Clg/yr ' ~
4oo
tect ----12"
,
.
.
_
C
.
~ '
base= 8
mW/m2_
_ _d
i! S
0-1 x (km) Fig. 6. Parameters: E* described in Fig. 3.
.
2_
=
E ; e*f / 3
,
0 (%)
1 0
500
~
9
%
%
1000
I
I
1
1500
x (kin)
B = 1.0, n = 3, r/b(Z) = NLO. Four time frames: t = (a) 0, (b) 150, (c) 500, and (d) 1000 My. Frames as
125
S.S. Shapiro et al./Lithos 48 (1999) 115-133
=
_
o My;
Vm, ax = 5 cndyr,
_ _=
z(II)
11,
_ : ~ ~
,
l
~,
~
9
i
I
i
~
[
= 1000 M y ; , V m a x = 3 cm/yr L__J
t
t
s
J
~,
.A
_
.,.
%
%
..~
~
~
,~
!
,_
_C(1000 ~ My)-C
(0),
T(IO00 My) - T(O)
" \ ~~ K ~ ~
0~'
V'- ~ /
-x /
x
x A; Az CBL
40
q
(g)
t"q
0
0
t(My) Fig. 7. Parameters as in Fig. 6. Frames as described in Fig. 4.
t(My)
126
S.S. Shapiro et al./Lithos 48 (1999) 115-133
the region of high stress is significantly reduced, as is the region of high viscosity (Fig. 4d). Subtracting the initial from the final composition (and temperature) fields illustrates which regions have gained or lost composition (and heated up or cooled down) and by how much (Fig. 4e). By plotting the area, A, of the - 0 . 1 composition difference contour (normalized to the combined area of the initial oceanic and continental CBLs) as a function of time (Fig. 4f), we see that most of the composition is removed from the tectosphere CBL within the first 50 My. After this time, the rate of removal diminishes dramatically, decreasing slowly to zero at about 600 My. After the initially weak (i.e., easily deformed) material is quickly washed away, the remnant continental CBL weakens gradually as a result of the combination of the hot material flowing along its base and the strong temperature-dependence of the material's viscosity. We interpret this temporal pattern as indicating the presence of two instabilities, a mechanical mode (t < 50 My) and a thermal ablation mode (t > 50 My). The change in the depth to the base of the CBL, A ZCBL-- ZCBL(0) --ZCBL(t), expressed as a percentage of the initial CBL thickness (ZcBL(0)), varies with time in a manner similar to A(t) (Fig. 4f). Such a correspondence is reassuring since both of these quantities were devised independently to estimate, as a function of time, the condition of the CBL. The conductive heat flux through the continental surface (Fig. 4g) increases by only a factor of about t w o (10 m W m -z ) throughout the calculation. The advective heat flux through the base of the domain is more variable. The intervals of large fluxes coincide with part of the tectosphere's TBL falling off and sinking through the base of the domain. Such occurrences are episodic. (Note that the time interval at which we stored the fields needed to compute regional fluxes is too coarse to fully resolve the peaks in the flux-time curve.) When we repeat the above experiment with B = 1, the results differ significantly. The CBL remains largely intact after 1000 My, although its initially horizontal base becomes rounded (Fig. 6). The function ~p(z)/p changes shape, gradually forming a profile which indicates that a cold, dense upper section of the tectosphere is being partially supported from below by a light CBL opposite to our initial
condition where the TBL extends beneath the CBL, causing the small positive deviation from zero in ~p(z)/p. This dense upper region is caused by the conductive cooling of the upper section of the tectosphere with composition effectively unable to diffuse, the isopycnic condition is no longer satisfied in this region. Throughout the experiment, both the average geoid height anomaly and the dynamic topography associated with the tectosphere CBL have geophysically reasonable values (Fig. 7a and b). Due to the addition of compositional buoyancy, the second invariant of the stress is reduced relative to the above experiment (B = 0). This lower stress yields a correspondingly higher viscosity field (Fig. 5c, Fig. 6b, Fig. 7c and Fig. 7d) which leads to a more stable CBL. The changes in the composition and temperature fields after 1000 My are confined to a much smaller area than in the previous experiment (Compare Fig. 4e and Fig. 7e). Again, the initially weak material is removed quickly from the tectosphere, although with the stabilizing effect of compositional buoyancy more material remains (Fig. 7f). Further, the higher viscosity due to the lower stress than in the B = 0 case inhibits the thermal ablation process so that it is barely observable in A(t). Here, A(t) and A ZcBL(t) diverge at about 200 My. After this time, A(t) is essentially constant whereas A ZCBL(t) decreases slightly, indicating that the CBL is thickening. From Fig. 6c and d, one can see that this thickening occurs in the center of the tectosphere CBL as a consequence of the cold "blobs" detaching from the tectosphere TBL. The plot of heat flux through the base of the domain (Fig. 7g) shows some of these instabilities. The heat flux through the top of the tectosphere, however, does not vary noticeably with time, holding at a value of about 20 mW m -2.
6. Bounds on E*, B,
and "qb(Z) from longevity
and geoid constraints To evaluate the relative effects of (1) activation energy, (2) compositional buoyancy, and (3) rheology on tectosphere stability and the resulting geoid signal, we plot the depth to the base of the CBL and the average geoid height anomaly over the continental tectosphere, at t = 1000 My, for each experiment
S.S. Shapiro et al./Lithos 48 (1999) 115-133
listed in Tables 1 and 2 (Fig. 8). We do not plot results for our experiments with E*/9, which are so
127
unstable that the tectosphere is removed in t << 1000 My, and the calculations terminated.
250
ill
i I
I
I D
I
6 N(m) Fig. 8. The depth to the base of the CBL (ZCBL) (ordinate) and the mean geoid height anomaly (gN) over the tectosphere CBL (abscissa) both at 1000 My. The depth to the base of the CBL is estimated from the median depth to the C = 0.1 contour within the tectosphere CBL. Triangles and squares indicate experiments with activation energies of E* = E* r and E*f/3, respectively. White, gray, and black symbols indicate experiments with buoyancy ratios of B = 0, 1, and 1.5, respectively. Small and large symbols indicate experiments with background viscosity profiles based on HGPA and NLO, respectively. The ordinates of the horizontal dashed lines indicate a (somewhat arbitrary) upper bound of ZCBL based on the inferred thickness of the continental tectosphere (Shapiro, 1995). The abscissas associated with the vertical dashed lines correspond to gNps _+ 2o- as determined by Shapiro et al. (1999). (a) Newtonian rheology. (b) Non-Newtonian rheology. Circles represent experiments with a background viscosity profile determined by NLO and an activation energy of E* -- E'f~ 1.25. White and gray symbols indicate experiments with buoyancy ratios of B = 0 and 1, respectively.
128
S.S. Shapiro et al./Lithos 48 (1999) 115-133
In both the Newtonian and non-Newtonian experiments, using an activation energy of 522 kJ mole -1, corresponding to dry olivine (Ashby and Verrall, 1977), assures stability regardless of the amount (within reasonable limits) of compositional buoyancy present. However, not all of these experiments produce geoid height anomalies consistent with those observed for platforms and shields (Shapiro et al., 1999). In fact, by comparing the average geoid height anomalies resulting from the E* = E~*f experiments (triangles in Fig. 8), one can see that simply satisfying the isopycnic hypothesis is insufficient to ensure a geoid height anomaly consistent with the observed geoid associated with platforms and shields one must also consider the dynamics of the flow. For example, the Newtonian experiments with initial B = 1.5, but not with B = 1, satisfy the geoid constraint. From the corresponding, more realistic, experiments with stress-dependent rheology, we estimate that the initial B is likely between about 0.9 and 1.3, depending on the background viscosity profile (Fig. 8b). (We refer to the "initial" B because, after 1000 My, both the temperature and composition are modified and the local value of B changes. For example, for the experiment shown in Fig. 6, at 1000 My, for z < 400 kin, 8p > 0 implies B < 1, while for z > 400 km, 8 p < 0 implies B > 1.) With an activation energy one third the above value (E~ef/3) and a stress-dependent rheology, compositional buoyancy plays a major role in stabilizing the continental tectosphere (compare the open with the shaded squares in Fig. 8b; compare also Figs. 3 and 6). After 1000 My, the CBLs with B = 1 are roughly 200 km thicker than those containing no compositional buoyancy. Compositional buoyancy has a similar effect on stability for the corresponding Newtonian experiments, except for those characterized with an activation energy of E ~*e f / 3 . In this case, B = 1 results in CBLs that are only about 50 km thicker than the corresponding CBLs that contain no compositional buoyancy. By comparing the Newtonian to the non-Newtonian experiments, we see that compositional buoyancy plays a larger stabilizing role with a non-Newtonian rheology. Compositional buoyancy reduces the stress within the boundary layer, which in conjunction with the stress dependence of viscosity, causes an increase in viscosity which stabilizes the CBL.
The two background viscosity models yield continental tectosphere thicknesses that are always within about 50 km of each other. In most, but not all cases, the model with a background viscosity described by NLO is the more stable one.
7. Discussion
The most important question our numerical experiments were designed to address is whether under any circumstances a thick, cold root beneath continents can be stable over time scales of a billion years or more. The alternative is that these roots are parts of massive convective downwellings (e.g., Pari and Peltier, 1996). We find that the effects of temperature dependent viscosity can stabilize these roots, with only slow ablation by small-scale convection. The wisps of cold material sinking into the deeper interior in our models are consistent with the seismic observations of Li et al. (1998), who found that beneath the deep root below the eastern United States there is little perturbation in the depth of the 400 km seismic discontinuity. They conclude that any cold downwellings in the transition zone in the region they studied must be weak ( < 50 K) and/or small-scale. Conclusions drawn from studies of boundary-layer dynamics depend strongly on assumptions concerning the temperature and stress dependence of viscosity and the magnitude of %, the typical value of background stress in the convecting mantle. From an analytical analysis of the stability of a constant viscosity continental tectosphere described by linear gradients in composition and temperature, assuming 77 = 10 21 Pa s, Stevenson (1979) found modes of instability with characteristic growth times as short as about 200 My w an order of magnitude less than that required by the above age constraints. If one considers the temperature dependence of viscosity and the fact that continents are cold, a constant viscosity of r/= 1021 Pa s is an unrealistically low estimate (e.g., Simons and Hager, 1997). An increase of only one order of magnitude would yield characteristic time constants comparable to the age of Earth. In addition, the strong temperature dependence of viscosity helps to confine the flow to a very
129
S.S. Shapiro et al./Lithos 48 (1999) 115-133
limited region, further stabilizing the tectosphere, as discussed below. In the high-stress regime, where zc >> 1 MPa (10 bars), Kincaid (1990) concluded that viscosity, not compositional buoyancy, is responsible for achieving long-term stability. Shapiro (1995) further demonstrated that, with zc 6 MPa (60 bars), 10-20 times that considered to be appropriate for most of Earth (e.g., Hager and O'Connell, 1981), a viscosity increase of a factor of about 20 between the TBL and the surrounding mantle is sufficient to maintain stability, regardless of the amount of compositional buoyancy (within reasonable bounds). These studies probably overestimate the role of tractions from mantle convection because, due to numerical constraints, they were carried out at a Rayleigh number lower than appropriate for Earth. Convective stresses scale as "ro=pgdaATRa~/3; since the value of AT assumed in these studies is a realistic one, higher Rayleigh numbers, corresponding to lower viscosities, result in thinner boundary layers, and lower stresses. Our transition stress value is based on a small grain size about one mm (Ashby and Verrall, 1977). As grain size increases, the transition stress
0 t = 1000 My; Vclip = 1 cm/yr
200-
decreases, so our choice of % = 1 MPa (10 bars) is a conservative one in the following sense: If we had assumed a larger grain size, the fluid within and beneath the tectosphere would have been more viscous, and the tectosphere would have been more likely to survive. In the low-stress regime, we find that destruction is achieved in two ways: (1) through a mechanical removal of material and (2) via a thermal ablation process in conjunction with the mechanical process. The initially weak regions of the tectosphere are washed away quickly by convective processes; the remaining material is removed more slowly. As the base warms through conduction, it becomes weaker due to viscosity's inverse dependence on temperature. Then, in its weakened state, it is swept away by the convection currents. To estimate the viscosity required to prevent this mechanical removal, we consider the viscosity corresponding with this "ablation front". For the experiments which yield CBLs that are stable over a billion year time scale (Fig. 8), the viscosity corresponding to the edge of the ablation front, where velocities become negligible, is approximately 10 20.5 Pa s (for a representative example, see Fig. 9). Once the viscosity of the tectosphere
--%_---i
\
r
i
\
~ ~ ~\..
- - ~
=
400
1500
1000
500
0
Fig. 9. Model parameters as in Figs. 6 and 7. Viscosity field (r/) (solid lines) with velocity arrows and the change in the composition field (see Fig. 7) (dashed lines) superposed; t = 1000 My. Velocity scaling is clipped at /"cut= 1 c m / y e a r to exhibit lower values more clearly. Contour intervals as in Fig. 4.
130
S.S. Shapiro et al./Lithos 48 (1999) 115-133
material is about a factor of ten greater than that of the underlying mantle, negligible flow occurs similar to the result from high-stress models (Shapiro, 1995). Further support for this interpretation comes from the work of Conrad and Molnar (1997). They used a semi-analytic approach to investigate RayleighTaylor-type instabilities of a more dense fluid layer of thickness, h, overlying a uniform halfspace. Within the layer, density increases linearly with distance above the boundary, as A p z/h, and viscosity varies exponentially, as r/0 exp(z/L). Their results show that flow is essentially confined to a layer of thickness z = 2L, corresponding to a depth range over which the viscosity varies by a factor of 10. As is the case with our numerical experiments, there is negligible flow excited in regions with higher viscosity. For a ratio L / h ~ 0.03, typical Rayleigh-Taylor growth times are of order 30 ~qo/ApgL. Assuming representative values of density contrast, A p = 100 kg/m3; tectospheric thickness, h = 300 km; reference viscosity, r/0 = 1 0 2~ Pa s; and viscosity e-folding distance, L = 10 km; the characteristic growth time is ~ 50 My. Such a short growth time is consistent with the results of our numerical experiments, which show a subtectospheric drip developing on a comparable time scale. Once this sublithospheric drip strips away the initial layer of thickness 2 L ~ 20 km, the gradient in viscosity at the base of the tectosphere is much sharper, resulting in a situation where L is of order 1 km. Since the growth time is inversely proportional to L, the rate at which the tectosphere ablates decreases by a comparable factor, with a characteristic growth time ~ 1 Gy. Fleitout and Yuen (1984) demonstrated that the combination of pressure- and temperature-dependent viscosity can help to stabilize a thick thermal boundary layer from convective disturbance. The temperature-dependence of viscosity stabilizes the cold (shallow) part of the TBL while the pressure-dependence of viscosity stabilizes the warmer (deeper) part. Our numerical representation of the continental tectosphere is stable even without the stabilizing effects of pressure-dependent viscosity. Doin et al. (1997) have extended these calculations, investigating the effects of rheology dependent on temperature, pressure, and composition, in a do-
main of width 2680 km and depth 670 km. They enforce a tectonic regionalization using velocity boundary conditions, with one third to one half of the surface, representing a continent, pinned to the surface, and the remaining domain driven by imposed horizontal velocities, representing oceanic plates. The velocity boundary conditions change with time such that the subduction zone is periodically adjacent to either edge of the tectosphere region, with one side or the other of the craton experiencing subduction every 90 My. They find that viscosity is the most important parameter in determining whether the tectosphere is stable, and that compositional buoyancy is not a primary contributor for the value of E* that they use, 310 kJ mole-1 (corresponding to 0.6 E'f). For models with strong pressure dependence of viscosity, there are two preferred lithospheric thicknesses, one corresponding to mature oceanic lithosphere and the other to the thickness of the subcontinental tectosphere. Doin et al. (1997) find somewhat higher rates of tectosphere ablation than do we. One reason is that the activation energy that they used, 310 kJ mole-l, is less than the value that we used, 522 kJ mole-l, the estimate for dry olivine (Ashby and Verrall, 1977). Also, their models have flow confined to the upper mantle, so there is strong shear between the base of their continental TBL and the bottom of their computational domain. Perhaps most importantly, their "continents" are narrow (900-1350 km wide) and are bounded by subduction zones at intervals of < 100 My. In contrast, on Earth, tectosphere is only rarely adjacent to a subduction zone. For example, Bostock (1998, Bostock, 1999) interprets the seismic structure of the Slave craton as resulting from a fossil subduction zone that has been preserved for > 1.9 Gy since the suturing of the Fort Simpson terrain and the Slave province. Forte et al. (1995) suggested that the buoyancy profile beneath continents reverses in sign at about 250 km depth. Specifically, Forte et al. (1995) proposed that the negative buoyancy in the upper 250 km of the subcontinental mantle is partially supported by underlying lighter material. Interestingly, some of our acceptable models exhibit this same buoyancy reversal (see, for example, Fig. 6). As we discussed above, this buoyancy reversal is caused by the conductive cooling of the tectosphere. This cool-
s.s. Shapiro et al.,/Lithos 48 (1999) 115-133
ing is dependent on the convecting system in which it is placed. In fact, the existence of a buoyancy reversal is very sensitive to several of our assumptions, with some acceptable models showing this reversal, and others not [see, for example, Fig. B.21 of Shapiro (1995)]. Hence, our study cannot be used to predict accurately whether such a reversal actually exists. For models with viscosity sufficiently high that the tectosphere survives, there is negligible deformation of the tectosphere. Thus, heat transport through the tectosphere in our models is controlled by conduction. The predicted mantle heat flow decreases from the suboceanic region to the interior of the tectosphere, as the depth to a given isotherm increases. Such variation is consistent with the careful analysis of heat flow observations by Jaupart and Mareschal (1999), who estimated variations in mantle heat flow and lithospheric thickness after stripping off the effects of near-surface heat production. They estimate that the conductive part of the subcratonic thermal boundary layer is 200-330 km thick. Only beneath that depth does advective heat transport become important. In contrast, Lenardic (1997) used the results of convection models that, like ours, include the effects of compositional buoyancy, to challenge this interpretation. In Lenardic's models, the viscosity of the material comprising the subcratonic regions is sufficiently low that recirculation within the crust and underlying buoyant material is driven by the convecting mantle. This recirculation leads to a substantial advective contribution to the heat flux. In his models, there is little variation in mantle heat flux between "cratonic" and surrounding regions. In our view, if the tectosphere is to survive, its viscosity must be so high that deformation is negligible; heat transport results from simple conduction, with negligible advective contribution. The results of Jaupart and Mareschal (1999), who explain observed heat flow variations of both short and long wavelengths in terms of observed crustal heat production and a conductive model of heat transport through the tectosphere, support our interpretation. In summary, the joint application of longevity and gravity constraints allows us to evaluate the importance of specific properties of a continental tectosphere in the low-stress regime. High viscosity is
131
crucial for the long-term survival of the tectosphere. Flow models characterized by the activation energy for dry olivine, 522 kJ mole-1, yield stable boundary layers that, once established, are stable, even with no compositional buoyancy present. However, activation energies, say tenfold smaller, are too low; they lead to a rapid (of order 10 My) destruction of the tectosphere. With an activation energy about 20% less than that estimated for olivine, temperature-dependent viscosity alone is sufficient to assure stability (Fig. 8). With lower values of activation energy, stability of an existing tectosphere can be achieved with the inclusion of compositional buoyancy. Compositional buoyancy plays a dual role within a thermal (and chemical) boundary layer: It (1) reduces the stress within the boundary layer and (2) counteracts the thermally-induced density increase. With a stress-dependent rheology, this reduction in stress results in an increase in viscosity which, in turn, inhibits a greater region of the boundary layer from deforming. Removal of volatiles by depletion would also increase the viscosity, providing a plausible mechanism contributing to the stabilization of the tectosphere (Pollack, 1986). If, for realistic activation energies, compositional buoyancy is not required to maintain a stable tectosphere, it is interesting to ask why the geoid observations (Shapiro et al., 1999) indicate that B ~- 1? If the tectosphere formed by advective thickening, the results of our numerical experiments provide a plausible answer. Formation via advective thickening requires that the material that now constitutes the rigid tectosphere was ductile enough to deform and thicken at the time of formation of the proto-tectosphere. Thus, the proto-tectosphere was likely to have been somewhat warmer than mature tectosphere is now in order that it had a sufficiently low viscosity to deform. Since it is the ratio of E * / T that governs viscosity, a higher T is equivalent to a lower E*. As our numerical experiments demonstrate, compositional buoyancy is required to stabilize the tectosphere at lower E*. Thus it seems likely that compositional buoyancy would also be important in stabilizing a somewhat hotter proto-tectosphere, formed under advective thickening, that later completely stabilized by moderate cooling. Such moderate cooling would be consistent with the estimate from the geoid signal associates with cratons
s.s. Shapiro et al./Lithos 48 (1999) 115-133
132 that the p r e s e n t - d a y
B is slightly less t h a n u n i t y
( S h a p i r o et al., 1999).
Acknowledgements W e t h a n k S c o t t D. K i n g for h e l p i n g us to i n c o r p o rate
chemical
variations
into the code,
ConMan,
P e t e r P u s t e r a n d M a r k S i m o n s for their c o n t r i b u t i o n s to o u r v e r s i o n o f ConMan, a n d C l i n t o n C o n r a d for discussions. Adrian Lenardic and Richard O ' C o n n e l l provided
insightful reviews.
This
work
was
sup-
p o r t e d b y N a t i o n a l S c i e n c e F o u n d a t i o n grants E A R 9506427 and C D A - 9 6 0 1 6 0 3 .
References Ashby, M.F., Verrall, R.A., 1977. Micromechanisms of flow and fracture, and their relevance to the rheology of the upper mantle. Philos. Trans. R. Soc. London A 288, 59-95. Bostock, M., 1998. Mantle stratigraphy and evolution of the Slave province. J. Geophys. Res. 103, 21183-21200. Bostock, M.G., 1999. Seismic imaging of lithospheric discontinuities and continental evolution. Lithos., this issue. Brooks, A.N., 1981. A Petrov-Galerkin finite element formulation for convection dominated flows. Ph.D. thesis, California Institute of Technology. Conrad, C.P., Molnar, P., 1997. The growth of Rayleigh-Taylortype instabilities in the lithosphere for various rheological and density structures. Geophys. J. Int. 129, 95-112. Doin, M.-P., Fleitout, L., Christensen, U., 1997. Mantle convection and stability of depleted and undepleted continental lithosphere. J. Geophys. Res. 102, 2771-2787. EkstriSm, G., Dziewonski, A.M., 1995. Improved models of upper mantle S velocity structure. EOS Trans. AGU 76, 421, (abs). Fleitout, L., Yuen, D.A., 1984. Steady state, secondary convection beneath lithospheric plates with temperature- and pressure-dependent viscosity. J. Geophys. Res. 89, 9227-9244. Forte, A.M., Dziewonski, A.M., O'Connell, R.J., 1995. Continent-ocean chemical heterogeneity in the mantle based on seismic tomography. Science 268, 386-388. Grand, S.P., 1994. Mantle shear structure beneath the Americas and surrounding oceans. J. Geophys. Res. 99, 11591-11621. Hager, B.H., 1991. Mantle viscosity: a comparison of models from postglacial rebound and from the geoid, plate driving forces, and advected heat flux. In: Sabadini, R., Lambeck, K., Boschi, E. (Eds.), Glacial Isostasy, Sea-Level and Mantle Rheology. Kluwer Academic Publishers, Dordrecht, 493-513. Hager, B.H., O'Connell, R.J., 1981. A simple global model of
plate dynamics and mantle convection. J. Geophys. Res. 86, 4843-4867. Jaupart, C., Mareschal, J.C., 1999. The thermal structure and thickness of continental roots. Lithos., this issue. Jordan, T.H., 1975. The continental tectosphere. Rev. Geophys. Space Phys. 13, 1-12. Jordan, T.H., 1978. Composition and development of the continental tectosphere. Nature 274, 544-548. Kincaid, C., 1990. The dynamical interaction between tectosphere and large scale mantle convection. EOS Trans. AGU 71, 1626, (abs). King, S.D., 1990. The interaction of subducting slabs and the 670 kilometer discontinuity. Ph.D. thesis, California Institute of Technology. King, S.D., Raefsky, A., Hager, B.H., 1990. ConMan: vectorizing a finite element code for incompressible two-dimensional convection in the earth's mantle. Phys. Earth Planet. Inter. 59, 195-207. Lenardic, A., 1997. On the heat flow variation from Archaen cratons to Proterozoic mobile belts. J. Geophys. Res. 102, 709-721. Li, A., Fischer, K.M., Wysession, M.E., Clarke, T.J., 1998. Mantle discontinuities and temperature under the North American continental keel. Nature 395, 160-163. Masters, T.G., Johnson, S., Laske, G., Bolton, H., 1996. A shear velocity model of the mantle. Philos. Trans. R. Soc. London, A 354, 1385-1414. Nakada, M., Lambeck, K., 1989. Late Pleistocene and Holocene sea-level change in the Australian region and mantle rheology. Geophys. J. Int. 96, 497-517. Pail, G., Peltier, W.R., 1996. The free-air gravity constraint on subcontinental mantle dynamics. J. Geophys. Res., 101, 28, 105-28, 132. Pearson, D.G., Shirey, S.B., Carlson, R.W., Boyd, F.R., Pokhilenko, N.P., Shimizu, N., 1995. Re-Os, Sm-Nd, and Rb-Sr isotope evidence for thick Archaean lithospheric mantle beneath the Siberian craton modified by multistage metasomatism. Geochim. Cosmochim. Acta 59, 959-977. Pollack, H.N., 1986. Cratonization and thermal evolution of the mantle. Earth Planet. Sci. Lett. 80, 175-182. Richardson, S.H., Gurney, J.J., Erlank, A.J., Harris, J.W., 1984. Origins of diamonds in old enriched mantle. Nature 310, 198-202. Shapiro, S.S., 1995. Structure and dynamics of the continental tectosphere. Ph.D. thesis, Massachusetts Institute of Technology. Shapiro, S.S., Hager, B.H., Jordan, T.H., 1999. The continental tectosphere and Earth's long-wavelength gravity field. Lithos., this issue. Simons, M., Hager, B.H., 1997. Localization of the gravity field and the signature of glacial rebound. Nature 390, 500-504. Stevenson, D.J., 1979. Double diffusive instabilities in the mantle. I.U.G.G. 17th Gen. Assembly, I.A.S.P.E.I. Canberra, Australian National University, 72 (abstr.). Su, W.-J., Woodward, R.L., Dziewonski, A.M., 1994. Degree 12 model of shear velocity heterogeneity in the mantle. J. Geophys. Res. 99, 6945-6980.
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Walker, R.J., Carlson, R.W., Shirey, S.B., Boyd, F.R., 1989. Os, Sr, Nd, and Pb isotope systematics of southern African peridotite xenoliths: Implications for the chemical evolution of subcontinental mantle. Geochim. Cosmochim. Acta 53, 15831595.
133
Wise, D.U., 1974. Continental margins, freeboard and the volumes of continents and oceans through geological time. In: Burk, C.A., Drake, C.L. (Eds.), The Geology of Continental Margins. Springer-Verlag Publishers, New York, 45-58.
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Lithos 48 (1999) 135-152
The continental tectosphere and Earth's long-wavelength gravity field Steven S. Shapiro 1, Bradford H. Hager * , Thomas H. Jordan Department of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA 02139, USA
Received 19 October 1998" received in revised form 8 February 1999; accepted 15 February 1999
Abstract To estimate the average density contrast associated with the continental tectosphere, we separately project the degree 2-36 non-hydrostatic geoid and free-air gravity anomalies onto several tectonic regionalizations. Because both the regionalizations and the geoid have distinctly red spectra, we do not use conventional statistical analysis, which is based on the assumption of white spectra. Rather, we utilize a Monte Carlo approach that incorporates the spectral properties of these fields. These simulations reveal that the undulations of Earth's geoid correlate with surface tectonics no better than they would were it randomly oriented with respect to the surface. However, our simulations indicate that free-air gravity anomalies correlate with surface tectonics better than almost 98% of our trials in which the free-air gravity anomalies were randomly oriented with respect to Earth's surface. The average geoid anomaly and free-air gravity anomaly over platforms and shields are significant at slightly better than the one-standard-deviation level: - 1 1 _+ 8 m and - 4 _+ 3 regal, respectively. After removing from the geoid estimated contributions associated with (1) a simple model of the continental crust and oceanic lithosphere, (2) the lower mantle, (3) subducted slabs, and (4) remnant glacial isostatic disequilibrium, we estimate a platform and shield signal of - 8 + 4 m. We conclude that there is little contribution of platforms and shields to the gravity field, consistent with their keels having small density contrasts. Using this estimate of the platform and shield signal, and previous estimates of upper-mantle shear-wave travel-time perturbations, we find that the average value of i?ln p/Oln vs within the 140-440 km depth range is 0.04 _ 0.02. A continental tectosphere with an isopycnic (equal-density) structure @ln p/Olnv s = 0) enforced by compositional variations is consistent with this result at the 2.0o" level. Without compositional buoyancy, the continental tectosphere would have an average 01n p / i ) l n v S ~ 0.25, exceeding our estimate by 10o-. 9 1999 Published by Elsevier Science B.V. All rights reserved. Keywords." Continental tectosphere" Earth; Long-wavelength gravity field; Geoid anomaly; Gravity anomaly
1. Introduction Motivated by seismological evidence (e.g., Sipkin and Jordan, 1975) and the lack of a strong correla-
* Corresponding author. ~Present address" Department of Physics, Guilford College, Greensboro, NC 27410, USA
tion between continents and the long-wavelength geoid (e.g., Kaula, 1967), Jordan (1975) proposed that continents are (1) characterized by thick ( ~ 400 km) thermal boundary layers (TBLs) which translate coherently during lateral plate motions, (2) stabilized against small-scale convective disruption by gradients in density due to compositional variations, and (3) not observable in the long-wavelength gravity
0024-4937/99/$ - see front matter 9 1999 Published by Elsevier Science B.V. All rights reserved. PII: S0024-4937(99)00027-4
136
S.S. Shapiro et a l . / Lithos 48 (1999) 135-152
field. The simple plate cooling model, which enjoys much success in describing the structure of oceanic TBLs, cannot be extended to explain thicker continental TBLs (Jordan, 1978). Instead, Jordan (1978) postulated that the thick continental TBL, continental tectosphere, was formed early in Earth's history by advective thickening and has been stabilized against convective cfisruption by the compositional buoyancy provided by a depletion of basaltic constituents. The isopycnic (equal-density) hypothesis (Jordan, 1988) predicts that the compositional and thermal effects on density cancel at every depth between the base of the mechanical boundary layer and the base of the TBL. Such a structure would be neutrally buoyant with respect to neighboring oceanic mantle, and would not be visible in the long-wavelength gravity field. There has been much discussion during the past 2 decades about the relations among the Earth's longwavelength gravity field, surface tectonics, and mantle convection. For example, there is an obvious association of long-wavelength geoid highs with subduction zones (Kaula, 1972; Chase, 1979; Crough and Jurdy, 1980; Hager, 1984) and with the distribution of hotspots (Chase, 1979; Crough and Jurdy, 1980; Richards and Hager, 1988). Most of the power in the longest wavelength geoid can be explained in terms of lower-mantle structure imaged by seismic tomography (e.g., Hager et al., 1985; Hager and Clayton, 1989; Forte et al., 1993a). This lower mantle seismic structure has been linked to tectonic processes, in particular, to the history of subduction (e.g., Richards and Engebretson, 1992). Although there is general agreement among geodynamicists that most of the geoid can be explained in terms of features such as subducted slabs and lower mantle structure, there is significant quantitative disagreement among the predictions of various models (e.g., Panasyuk, 1998). Thus, it is not possible to estimate with high confidence the "residual geoid" not explained by lower mantle structure. The contribution to the geoid of upper-mantle structures, including variations in the thickness of the crust and lithosphere, is a question whose answer is still disputed. Assuming that plates approach an asymptotic thickness of approximately 120 km after cooling about 80 My, the geoid would be expected to be higher by roughly l0 m over continents and over
midoceanic ridges than over old ocean basins due to the density dipole associated with isostatic compensation (Haxby and Turcotte, 1978; Parsons and Richter, 1980; Hager, 1983). At intermediate to short wavelengths, the expected changes in the geoid over these features are observed (e.g., Haxby and Turcotte, 1978; Doin et al., 1996), but the isolation of the geoid signatures of these features at long wavelengths is problematic. Using broad spatial averages over selected areas, Turcotte and McAdoo (1979) concluded that there is no systematic difference in the geoid signal between oceanic and continental regions. But, Souriau and Souriau (1983) demonstrated that there is a significant correlation between the geoid (spherical harmonic degrees l = 3-12) and the tectonic regionalization of Okal (1977). From degree-by-degree correlations (l = 2-20), Richards and Hager (1988) observed a weak association between geoid lows and shields. On the other hand, Forte et al. (1995) reported that the degree 2-8 geoid correlates significantly (99% confidence) with an ocean-continent function. Were there a significant ocean-continent signal, the continental tectosphere might have a substantial density anomaly associated with it, and might therefore be expected to play an active role in the largescale structure of mantle convection. For example, Forte et al. (1993b) and Pari and Peltier (1996), in their preferred models, assumed linear relationships between seismic velocity anomalies and density anomalies. They proposed dynamic models of the long-wavelength geoid in which the high velocity roots beneath continents are cold, dense downwellings in the convecting mantle. Such downwellings would depress the surface of continents dynamically by about 2 km (Forte et al., 1993b). The lack of significant temporal variation in continental freeboard over geologic time would require that these convecting downwellings be extremely long-lived and translate coherently with the continents (e.g., Gurnis, 1993). On the other hand, Hager and Richards (1989) and Forte et al. (1993b) found the best fits of their dynamic models to the geoid by assuming an unusually small global proportionality between seismic velocity anomalies and density anomalies in the upper mantle. Forte et al. (1995) showed that they could improve their fit to the geoid if they allowed subcontinental regions to have a different proportion-
137
S.S. Shapiro et al./Lithos 48 (1999) 135-152
Table 1 GTR1 (Jordan, 1981) Region Definition Oceans A B C Continents Q P S
1996), referred to the hydrostatic figure of Earth (Nakiboglu, 1982) (Fig. 2a). Although we use GTR1 (and coarser regionalizations created by combining some of these regions) for the bulk of this study, we also compare our results with those obtained using the tectonic regionalizations of Mauk (1977) and Okal (1977), as well as the ocean-continent function. Because the geoid spectrum is red, with the rootmean-square (rms) value of a coefficient of degree 1 decreasing roughly as 1-2 , and because the longest wavelengths are likely dominated by the effects of density contrasts in the lower mantle (Hager et al., 1985), we also investigate the relationship between GTR1 and free-air gravity anomalies. The gravity field at spherical harmonic degree 1 is proportional to 1-1, so the gravity anomalies are expected to have correspondingly smaller long-wavelength variations than the geoid does. We calculate regional averages of the geoid and the gravity field and estimate their uncertainties. Further, we try to refine the estimate of the contribution of the continental tectosphere to the geoid by
Fractional area (%)
Young oceans (0-25 My) Intermediate-age oceans (25-100 My) Old oceans ( > 100 My) Phanerozoic orogenic zones Phanerozoic platforms Precambrian shields and platforms
61 13 35 13 39 22 10 7
ality constant between velocity and density anomalies beneath continents than beneath oceans. To quantify the association of surface tectonics and Earth's gravity field, we investigate the significance of the association between the six-region global tectonic regionalization GTR1 (Jordan, 1981) (Table 1, Fig. 1) and the geoid, EGM96 (Lemoine et al., i
I
A
B
C
Q
P
I
S
Fig. 1. Tectonic regionalization, GTR1 displayedusing a Hammerequal-area projection. See Table 1 for a description of each region.
S.S. Shapiro et al./Lithos 48 (1999) 135-152
138
I
o!
il
~
-lO0 -80
' ........
~-'"';':.,.~:;~,
-60
' . . . . . . ' .......
-40
" ..................
-20
0
Meters
~r;;~,
20
~'~
:'-
40
. . . . . . . . . . . . . . . .
60
80
lO0
139
s.s. Shapiro et al./Lithos 48 (1999) 135-152
subtracting other contributions from the geoid estimates. By combining the upper-mantle shear-wave travel-time anomalies associated with platforms and shields (Shapiro, 1995) and the results from this study, we estimate, with uncertainties, the average of aln p/01n vs within the depth range 140-440 km, and compare our estimate with the isopycnic hypothesis of Jordan (1988).
Table 2 Okal (1977) Region
Definition
Fractional area (%)
D C B A T M S
Ocean (0-30 My) Ocean (30-80 My) Ocean (80-135 My) Ocean ( > 135 My) Trenches and marginal seas Phanerozoic mountains Shields
12.0 30.1 12.3 2.5 10.9 11.6 20.4
2. Tectonic regionalization and inversion GTR1 and regionalizations published by Okal (1977) (Table 2) and Mauk (1977) (Table 3) contain six, seven, and 20 regions, respectively. Both GTR1 and the regionalization of Mauk (1977) are defined on a grid of 5 ~ • 5 ~ cells, whereas the model of Okal (1977) is defined using 15 ~ 2 1 5 ~ and 10~ 2 1 5 ~ cells. The regionalization of Mauk (1977) allows for as many as 10 regions to be represented in a given cell, while the other regionalizations are defined with only one region per cell. In GTR1, the three oceanic regions (including marginal basins) are defined by equal increments in the square root of crustal age: 0-25 My (A), 25-100 My (B), and > 100 My (C) and the continental regions are classified by their generalized tectonic behavior during the Phanerozoic: Phanerozoic orogenic zones (Q), Phanerozoic platforms (P), and Precambrian shields and platforms (S). Like GTR1, the oceanic regions of Mauk (1977) are based largely on crustal age. However, the continental regions of Mauk (1977) are classified by age rather than by their tectonic behavior. The more complex parameterization associated with the regionalization of Mauk (1977) does not offer us any significant advantage over GTR1; as we show through representative projections, the platform and shield signatures from the regionalization of Mauk (1977) and from GTR1 are consistent with each other and only significant at slightly better than the one-standard-deviation level. The regionalization of
Okal (1977) is limited in the accuracy of its designation of regions. For example, Okal (1977) labels the entire continent of Antarctica a shield, whereas a significant fraction (-~ 1 / 3 ) is orogenic in nature. Okal (1977) also classifies some islands (e.g., Iceland and Great Britain) as shields. Misidentifications such as these might have a significant effect on results from associated data projections. In general, a tectonic regionalization containing N distinct regions can be described by N functions, R n ( n = 1, N ) , each having unit value over its region and zero elsewhere. By combining regions, we can construct other, coarser regionalizations. For example, by consolidating young oceans (A), intermediate-age oceans (B), and old oceans (C) of GTR1, into one region, and Q, P, and S, into another region, we can create a two-component (ocean-continent) tectonic regionalization (ABC, QPS). For much of this analysis, we combine regions P and S into one region (PS). For any such regionalization, we expand each R n in spherical harmonics, omitting degrees zero and one from our analysis because geoid anomalies are referred to the center of mass and any rearrangement of mass from internal forces cannot change an object's center of mass. With coefficient R nlm representing the ( l , m ) harmonic of region n, and coefficient d Im representing the ( l , m ) harmonic of the observed
Fig. 2. (a) Geoid, l = 2-36 (EGM96; Lemoine et al., 1996), referred to the hydrostatic figure of Earth (Nakiboglu, 1982); (b) Projection of (a) onto (A, B, C, Q, PS), and (c) Residual: (a-b). All plots are displayed using a Hammer equal-area projection with coastlines drawn in white. Negative contour lines are dashed and the zero contour line is thick. The contour interval is 10 m.
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S.S. Shapiro et a l . / Lithos 48 (1999) 135-152
Table 3 Mauk (1977) Region Oceans 1 2 3 4 5
Definition
7
Anomaly 0-5 (0-10 My) Anomaly 5 - 6 (10-20 My) Anomaly 6-13 (20-38 My) Anomaly 13-25 (38-63 My) Late Cretaceous sea floor (63-100 My) Early Cretaceous sea floor (100-140 My) Sea floor older than 140 My
Continents 8 9 10 11 12 13 14 15 16 17 18 19 20
Island arcs Shelf sediments Intermontane basin fill Mesozoic volcanics Cenozoic volcanics Cenozoic folding Mesozoic orogeny Post-Precambrian undeformed Late Paleozoic orogeny Early Paleozoic orogeny Precambrian undeformed Proterozoic shield Archaean shield
6
Fractional area (%) 61.5 4.0 10.4 6.9 10.2 21.1 5.4 3.5 38.4 1.4 7.1 0.7 0.4 1.4 1.8 2.7 9.5 1.9 1.8 1.5 6.2 2.0
(or model) geoid or gravity field, we use a leastsquares approach to solve: e/nm~n = d im
(l)
(summation convention implied here and below) for the regional averages, %. We include the additional constraint: An'Yn = 0
(2)
where A n represents the surface area spanned by region n. This constraint ensures that % have a zero (weighted) average, as, by definition, do the geoidheight (and free-air gravity) anomalies. The weighted-least-squares solution can be written:
y : [R r WR]-1Rr Wd
We next consider the effect of errors in d on our analysis. Although we have available the covafiance matrix for EGM96, this weight matrix is not the appropriate one for our analysis. As discussed previously, most of the power in the long wavelength parts of the geoid is the result not of surface tectonics, but of deep intemal processes. Unfortunately, the contribution of these deep processes cannot be determined to anywhere near the accuracy of the observed gravity field, so the covariance matrix will be swamped by the contributions of the errors due to neglecting important dynamic processes. Quantitative estimation of the errors associated with estimates of the contributions of these deep processes has rarely been attempted (Panasyuk (1998) is an exception). Here, we simply assume the identity matrix as our default weight matrix. For this matrix, the relative error in the harmonic expansion of the geoid increases as 12 (or as l for the gravity anomalies). This behavior is qualitatively consistent with the result that dynamic models of the geoid do better at fitting the longest wavelength components and progressively worse at fitting shorter wavelength components, for example, because the effects of lateral variations in viscosity become more important at shorter wavelengths (e.g., Richards and Hager, 1989). The sole exception to the identity weight matrix is our application of a large weight, 1000, to the surface-area constraint. Results from our inversions are insensitive to the value of this weight, so long as it is not less than ten times the weight associated with the data (in our case unity) nor so large ( > 106 times the data weight) that the inversion becomes numerically unstable.
(3)
where the values R/nm and A n are the elements of the matrix R, W is a weight matrix constructed from the covariance matrix associated with d, Yn are the elements of the vector y, and dtm and zero constitute the vector d.
3. Statistical analysis procedure Because neither the geoid nor the regionalization have white spectra, we do not use common statistical estimates of uncertainties. In fact, their spectra are quite red, implying that uncertainties in parameter estimates based on the assumption of white spectra will be substantially smaller than the actual uncertainties. Through the use of Monte Carlo techniques, we incorporate the spectral properties of these fields in our estimates of parameter uncertainties. For each
141
S.S. Shapiro et al./Lithos 48 (1999) 135-152
of 10,000 trials, we (1) randomly select an Euler angle triple from a parent distribution in which all orientations are equally probable and then, in accord
with the selected triple, rigidly rotate the sphere on which the data residuals (dies = d lm - R t m. y.) are defined, with respect to the sphere on which the surface
1000 0 -50
0
50
-50
0
50
-50
(m)
0
50
0
50
(m)
3000 2000
1000 _.,i
0 -50
(m)
(m)
3000 2000
1000 0 -50
0
50
0
5
10
15
20
Fig. 3. Non-hydrostatic geoid (EGM96, l = 2-36): Histograms of parameter values (a) 7A, (b) YB, (C) YC, (d) ~/Q, (e) YPs obtained from projections onto the tectonic sphere of the correlated data combined with 10,000 random orientations of the data residual sphere, characterized by a~m (see text). Gaussian distributions, determined by the standard deviation, mean, and area of each histogram, are superposed. (f) Histogram of variance reduction resulting from 10,000 random rotations of the data sphere, characterized by dtm, with respect to the tectonic sphere. The shaded and unshaded arrows indicate the variance reductions associated with the actual orientation and the maximum variance reduction, respectively.
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S.S. Shapiro et al. / Lithos 48 (1999) 135-152
tectonics are defined ("tectonic sphere"), (2) com71m bine the rotated data residuals dre s ( " ~ " denotes rotated) with the correlated data to produce pseudo
data, ( d lm -" dre 7lms nt- Rnlm")/n)' and (3) project d Im onto (A, B, C, Q, PS). The resulting histograms of parameter values (e.g., YA, TB, YC. . . . ) approximate
1500 1000 500 0 -10
0
10
-10
0
10
0
10
-10
0
10
2000 1500 1000 500 0 -10
(regal) ,
2000
|
,
PS
1500 1000 500 0 -10
0
10
0
5
10
Fig. 4. Free-air gravity (l = 2-36): Histograms of parameter values (a) YA, (b) YB, (c) Yc, (d) ye, (e) "~PS obtained from projections onto the tectonic sphere of the correlated data combined with 10,000 random orientations of the data residual sphere, characterized by d"Tin (see text). Gaussian distributions, determined by the standard deviation, mean, and area of each histogram, are superposed. (f) Histogram of variance reduction resulting from 10,000 random rotations of the data sphere, characterized by dtm, with respect to the tectonic sphere. The shaded and unshaded arrows indicate the variance reductions associated with the actual orientation and the maximum variance reduction, respectively.
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Gaussian distributions and, because the correlated signal is added to the rotated data residual before projecting the composite, the resulting histograms of parameter values are centered approximately on the parameter values corresponding to the actual orientation of the "data sphere" with respect to the tectonic sphere (Figs. 3 and 4). We take these latter parameter values as our parameter estimates and the standard deviations of these approximately Gaussian distributions as the parameter uncertainties. Alternatively, we could assign random (white noise) values to each coefficient describing the data-residual sphere while constraining its power spectrum to be unchanged through a degree-by-degree scaling. Histograms resulting from this approach yield very similar distributions and virtually the same values for the parameter estimates and their standard errors (Shapiro, 1995). If one relaxes the constraint by requiring only that the total power remains unchanged, then the resulting histogram distributions are narrower than the corresponding ones for which the spectra were scaled degree-by-degree. These smaller values for the standard errors in the parameter estimates likely coincide (Shapiro, 1995) in the limit of large numbers of trials with those determined from the elements of the variance vector V ~ /~'2ost diag{[R T WR]-I}, where Xpost, : is the (postfit) X-V per degree of freedom. As a criterion for the success of the model in fitting the data, we use the percent fractional difference in the prefit and postfit X 2. This percent variance reduction associated with each projection, i.e., inversion, is thus defined by 100[1-(X2o~t/X2re)]. From the results of the random rotations of the data sphere with respect to the tectonic sphere, we estimate significance levels in the variance reduction associated with each projection. Specifically, we associate the fraction of trials that yield lower variance reductions than the actual orientation with the confidence level of the variance reduction.
4. Projections Table 4 shows the regional averages and their corresponding statistical standard errors obtained by separately projecting the geoid and the free-air gravity anomalies onto (A, B, C, Q, PS). Fig. 3a-eFig. 4 a - e graphically display the 10,000 parameter estimates obtained from the Monte Carlo simulations that lead to the uncertainties given in Table 4. With the geoid, only regions (C) and (PS) have averages which are larger than their standard errors. However, the significance of these averages is only slightly above the one-standard-deviation level. For example, with 95% (2 o-) confidence, the geoid signature associated with platforms and shields is in the range - 2 7 to + 6 m, a rather broad range which does not even significantly constrain the sign of this signal. The projection of the geoid onto (A, B, C, Q, PS) is shown in Fig. 2b and further demonstrates that very little of the long-wavelength non-hydrostatic geoid can be explained simply in terms of surface tectonics. The magnitude of the geoid signal that is uncorrelated with (A, B, C, Q, PS) (Fig. 2c) is essentially the same as that of the geoid anomalies themselves, given by EGM96. Using the free-air gravity yields a somewhat different result: four regions have averages larger than their standard errors (Fig. 4a-e). The significance of three of these averages is at or below the 1.5o- level and the significance of the fourth, TQ, is at the 2.5o- level (Table 4). With 95% confidence, the free-air gravity signature associated with platforms and shields is in the range - 1 0 to + 1.4 mgal. Like with the geoid, this range is rather large and does not significantly constrain the sign of this signal. However, unlike the geoid projection, which explains less of the variance than about two-thirds of the random orientations of the data sphere (Fig. 3f), the free-air gravity projection explains more of the variance than about 98% of projections corresponding with random orientations
Table 4 EGM96 (l = 2-36): Regional averages and statistical standard errors from projections of the geoid and of perturbations to the free-air
gravity onto (A, B, C, Q, PS) Geoid (m) Gravity (mgal)
TA
TB
"Yc
TQ
]/ps
0 _+ 12 4.0 _+3.4
1.8 + 5.6 - 1.7 + 1.7
17 _+ 16 -4.8 +_3.9
-4.8 _+11 6.4 _+2.5
- 10.5 _+8.3 -4.3 _+2.9
S.S. Shapiro et al./Lithos 48 (1999) 135-152
144
of the data sphere (Fig. 4f). However, this reduction in variance is only about 6% and does not produce an impressive fit. Interestingly, Monte Carlo simulations using degrees 2-12 yield confidence levels of less than 30%, suggesting that the association between free-air gravity anomalies and surface tectonics is stronger in the higher frequencies. Using the regionalization of Mauk (1977) (the full 20-region tectonic sphere as well as some representative groupings of these regions) leads to results similar to those obtained from GTR1. In no case do we find a significant signal that can be linked with the continental tectosphere. Combining the regions of Mauk (1977) into three groups based on crustal age (regions [1-7], [8-14, 16-17], [15, 18-20]), yields regional averages which are roughly the same magnitude as their corresponding uncertainties (Table 5) and a variance reduction of about 6%. Another continental grouping ([1-7], [8-10, 12-13], [11, 14-
Table 5 EGM96 (1 = 2-36): Regional averages and statistical standard errors from projections of the geoid onto several regionalizations based on Mauk (1977). Group 1: ([1-7], [8-14, 16-17], [15, 18-20]); Group 2: ([1-7], [8-10, 12-13], [11, 14-17], [18-20]); Group 3 : 2 0 separate regions Region
Oceans 1 2 3 4 5 6 7 Continents 8 9 10 11 12 13 14 15 16 17 18 19 20
Group 1 3, (m)
Group 2
Group 3
3' (m)
3' (m)
9• 9• 9• 9• 9• 9• 9•
9• 9• 9• 9• 9• 9• 9•
-13• 10• 9• 5• 4• 15• 40•
- 22 + 16 - 2 2 + 16 - 2 2 + 16 - 2 2 + 16 - 2 2 + 16 - 2 2 + 16 - 2 2 + 16 -7+9 -22+16 - 2 2 + 16 -7 +9 - 7+ 9 -7 +9
- 9 + 15 - 9 + 15 - 9 + 15 - 19+ 14 - 9 + 15 - 9 + 15 - 19+ 14 -19+14 -19+14 - 1 9 + 14 - 12 + 13 - 12 + 13 - 12+ 13
127 + 4 6 - 4 1 + 19 62+44 - 10+57 27+41 6+37 -34__+ 22 4+14 -61+33 -60+26 37___ 28 - 27 + 15 -23+23
17], [18-20]) based instead on a combination of age and tectonic behavior, yields similar (insignificant) results (Table 5), and even produces a slightly smaller variance reduction than the previous model, which was based on one fewer parameter. On the other hand, when one uses the full 20-region tectonic sphere, the variance reduction associated with the projection of the data sphere is about 20%. This result by itself is not particularly surprising since one would expect the variance reduction to increase with the number of model parameters. However, using this regionalization, less than 10% of our Monte Carlo simulations result in a greater reduction in variance. While this result does not allow us to reject a strong association between the geoid and the surface tectonics defined by Mauk (1977), the large relative uncertainties (and even differences in sign) associated with old continents (Table 5) suggest that this association is indeed weak. In addition, there are only three regions (8, 9, and 17) that have average values that differ from zero by more than 2 o-. Although there is substantial uncertainty in the predictions of models of the contribution of other processes to the long-wavelength geoid, perhaps we could better isolate the tectosphere's contribution by subtracting from the observed (non-hydrostatic) geoid the effects of previously modeled components" (1) a simplified representation of the upper 120 km based on the oceanic plate cooling model and a uniform 35-km-thick continental crust ( l = 2-20) (Hager, 1983); (2) the lower mantle ( l = 2-4) (Hager and Clayton, 1989); (3) slabs (l = 2-9) (Hager and Clayton, 1989); and (4) remnant glacial isostatic disequilibrium (l = 2-36) (Simons and Hager, 1997). Separately projecting each of these four contributions to the model geoid onto (A, B, C, Q, PS) yields the results given in Tables 6 and 7. Our resulting model (residual) geoid, TECT-1 (Fig. 5), provides an estimate of the contributions to the geoid of the upper mantle structure below 120 km depth, excluding subducted slabs. For TECT-1, Yr,s = -8 _ 4 m (Table 6). The projections of TECT-1 separately onto (A, B, C, Q, PS), (ABC, QPS), and (ABCQ, PS) lead to reductions in variance that are listed in Table 7. From the percent of random trials that yield smaller variance reduction than that of the actual orientation (confidence level), it is clear that the geoid signal
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S.S. Shapiro et al./Lithos 48 (1999) 135-152
Table 6 Regional averages and statistical standard errors from projections onto (A, B, C, Q, PS), corresponding to contributions to the geoid from five model geoids - - each representing a separate contribution to the geoid. The bottom two represent projections of TECT-1, separately, onto (ABC, QPS) and (ABCQ, PS) Geoid contributors
TA (m)
YB (m)
7c (m)
TQ (m)
")/PS (m)
Upper 120 km Lower Mantle Slabs Post-Glacial Rebound TECT-1
4.3 • 1.0 - 5 +_ 32 - 11 _+9 1 _+0.5 3+ 5
- 3.1 _+0.5 20 + 19 - 5 _+4 1 _+0.3 2+ 3
- 6.1 + 1.0 35 +_ 34 - 1 + 11 0.7 + 0.5 4_+6
3.1 + 0.9 - 76 _+48 21 _+ 10 -0.2 _+0.3 - 1 _+4
3.5 _+0.8 34 _+ 35 -7_+ 6 - 3 _+0.5 - 8 _+4
TECT-1/(ABC, QPS) TECT- 1/(ABCQ, PS)
2.5 + 2 1.6 + 0.8
2.5 ___2 1.6 ___0.8
2.5 + 2 1.6 + 0.8
-4 + 3 1.6 + 0.8
-4 • 3 - 8+ 4
At these w a v e l e n g t h s (l = 2 - 3 6 ) , if there were no contribution from density contrasts at depths greater than 120 km, the geoid a n o m a l y associated with isostatically c o m p e n s a t e d platforms and shields w o u l d be about + 10 m, referenced to old ocean basins, or 0 m, referenced to ocean crust of zero age or to y o u n g continental crust (e.g., Hager, 1983). Our estimate of the geoid a n o m a l y associated with old ocean basins, from the TECT-1 projections, is 4 + 6 m, for oceans 0 - 2 5 M a is 3 + 5 m, and for y o u n g continents is - 1 +__4 m. D e p e n d i n g on w h e t h e r we take old oceans, y o u n g oceans, or y o u n g continents as the reference value, our estimate of the signal due to the tectosphere alone, correcting for the effects of the crust, w o u l d be - 2 2 m, - 1 1 m, or - 7 m. B e c a u s e the old oceanic regions m a y still have some residual effect of subduction included in their estimate, and because the area-weighted average of y o u n g oceans and y o u n g continents is close to
represented by TECT-1 is, a m o n g these choices, best r e p r e s e n t e d by the t w o - r e g i o n regionalization: ( A B C Q , PS). A l t h o u g h the projection of TECT-1 onto ( A B C Q , PS) results in a variance reduction of only about 3%, this value exceeds those obtained from almost 95% of the projections associated with r a n d o m rotations of the data sphere. This result is consistent with the r o u g h l y 20- result associated with the platform and shield signal represented in TECT-1 (Table 6), but contrasts m a r k e d l y with the results for the five-region grouping (A, B, C, Q, PS), where the actual orientation of the data sphere explains m o r e of the variance than only 54% of the r a n d o m orientations. This apparent discrepancy arises because r a n d o m orientations of the other tectonic regions can " l o c k o n " to regional features in the geoid such as those associated with subduction zones, providing a better fit to the synthetic geoids globally, but not in regions spanned by the projection of PS.
Table 7 Variance reductions and the corresponding confidence levels associated with the projection onto different groups of tectonic regions of five model geoids - - each representing a separate contribution to the geoid. Confidence level represents the percent of random trials that yield a smaller reduction in variance than that of the actual orientation of each geoid contributor Geoid Contributor
Projection
Variance reduction (%)
Confidence (%)
Upper 120 km Lower Mantle Slabs Post-Glacial Rebound TECT-1
A, B, C, Q, PS A, B, C, Q, PS A, B, C, Q, PS A, B, C, Q, PS A, B, C, Q, PS
80 37 13 19 4
100 88 77 100 54
TECT- 1 TECT- 1
ABC, QPS ABCQ, PS
2 3
78 94
S.S. Shapiro et al./ Lithos 48 (1999) 135-152
146
-100 -80
-60
-40
-20
0
Meters
20
40
60
80
100
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s.s. Shapiro et al./ Lithos 48 (1999) 135-152
zero, we retain the estimate of - 8 due to the continental tectosphere.
m as the signal
5. Estimate of ~ In p / ~ lnv s
The isostatic geoid height anomaly, g N, associated with static density anomalies can be calculated for each lateral location from (e.g., Haxby and Turcotte, 1978): -27rG gU=
~
g
f A p( z) zdz
(4)
where G is the universal gravitational constant, g is the acceleration due to gravity, and A p ( z ) is the anomalous density at depth z. The integration extends from the surface to the assumed depth of compensation. Assuming that ~ln p/i31n us is constant within a specified depth interval, we may write the scaling there between fractional perturbations in density and shear-wave velocity as:
~
~-
Olnvs
~
1
Table 8 S12_WM13 (Su et al., 1994) (/= 1-12): Platform and shield averages and uncertainties corresponding to one-way S-wave travel-time anomalies (Shapiro, 1995) Depth interval (km)
(AT/T)PS (%)
140-240 240-340 340-440
- 2.3 _ 0.2 - 1.6 _ 0.2 - 1.0 + 0.2
as the sum of the anomalies for these layers. Using the travel-time perturbations (Shapiro, 1995) (Table 8) and g Nps = Yl~S ~ - 7 . 7 + 3.9 m, we find that for platforms and shields,the average value of Oln p / O l n v s is about 0.041 ___0.021. (This estimate of standard error is based only on that of g Nps. The uncertainties associated with the regionally averaged travel-time perturbations have a much smaller effect on the value of Oln p/Olnt, s than the uncertainty associated with the geoid and are therefore ignored.)
6. Discussion
(5)
where } is obtained, for example, from the radial earth model PREM (Dziewonski and Anderson, 1981), and the fractional perturbations in shear wave velocity A vs/-~ s are equal to the negative of the fractional travel-time perturbations A~-/u for small perturbations. We base the subsequent calculation on a depth of compensation of 440 km. Below this depth, we assume that there is no platform and shield contribution to the geoid, as there is no significant distinction at such depths between the shear-wave signal beneath platforms and shields and the global average (Shapiro, 1995). Using S12_WM13 (Su et al., 1994), we calculated regional averages of one-way shear-wave travel-time perturbations for 100-km-thick layers between 140 and 440 km depth. We then approximate the integral of the depth-dependent density anomaly
None of the projections based on (1) the non-hydrostatic geoid, (2) free-air gravity anomalies, or (3) our model geoid, TECT-1, yields a platform and shield signal that is significant at a level exceeding about 2.00-. Our conclusion is in accord with that reached by Doin et al. (1996) using a geologic regionalization based on the tectonic map of Sclater et al. (1980). They estimated that shields have a geoid difference from midoceanic ridges of between - 1 0 m and 0 m; their corresponding estimate for platforms, which they keep as a separate region, is - 4 m to 1 m, while they found essentially no difference in geoid for tectonically active continental areas and ridges. They were unable to estimate formal errors because of the previously discussed red nature of the spectra, but these values represent their subjective estimates of confidence intervals. Although there are many differences in detail between
Fig. 5. (a) TECT-1, l = 2-36; (b) Projection of (a) onto (A, B, C, Q, PS), and (c) Residual: (a-b). All plots are displayed using a Hammer equal-area projection with coastlines drawn in white. Negative contour lines are dashed and the zero contour line is thick. The contour interval is 10 m.
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S.S. Shapiro et al./Lithos 48 (1999) 135-152
their study and ours, their estimates fall within our uncertainties, and their conclusion that the tectosphere is compositionally distinct is consistent with ours. These observations differ substantially from the highly significant (99% confidence) correlation, reported by Forte et al. (1995), between an ocean-continent function and the non-hydrostatic long-wavelength ( / = 2 - 8 ) geoid. However, a correlation coefficient ( r ) between different fields defined on a sphere is only meaningful (subject to tests of significance) for fields with (significantly) non-white spectra if correlation coefficients are determined separately for each spherical harmonic degree of interest (Eckhardt, 1984). Given the appropriate number of degrees of freedom associated with the correlation, one can nonetheless estimate the confidence level corresponding to the assumption that the true correlation is zero. Therefore, we estimate the effective number of degrees of freedom in the analysis of Forte et al. (1995) and, using this value, estimate the probability that the correlation which they obtained is significantly different from zero. Under the conditions outlined above, we can estimate the effective number of degrees of freedom using Student's t distribution. For uncorrelated fields, the quantity t = r[ v / ( 1 -/.2)]1/2 can be described by Student's t distribution with v degrees of freedom (e.g., Cramer, 1946; see also O'Connell, 1971). We create 10,000 degree-eight fields, each with the same spectral properties as the non-hydrostatic geoid, by randomly selecting coefficients from a uniform distribution and then scaling them degree-by-degree so that the power spectrum of each "synthetic" field matches that of the geoid. From these synthetic fields and an ocean-continent function derived from GTR1, we generate a collection of 10,000 correlation coefficients (Fig. 6a). We then estimate v by minimizing the X 2 in the fit of Student's distribution to this set of correlation coefficients (Fig. 6b). Fig. 6c,d,e
demonstrate the sensitivity of the fits to the value of v. As shown, values of v which differ from the estimated value (v = 30) by even 5 degrees of freedom, noticeably degrade the fit. The correlation coefficient corresponding to the geoid and (ABC, QPS) (l = 2 - 8 ) is - 0.18. However, using the geoid and an ocean-continent function (1 = 2-8) derived from the 5 ~ 2 155~ tectonic regionalization of Mauk (1977), we obtained the same value ( - 0 . 2 8 ) as Forte et al. (1995). With the regionalization of Mauk (1977), simulations like those described above yield 31 as the estimate of the effective number of degrees of freedom. The significance levels of the correlations associated with the GTR1 and Mauk (1977) ocean-continent functions are, respectively, about 85% and 95%. The dominant degree-two term in the geoid governs this correlation and highlights a difficulty associated with attaching significance to the correlations between such fields. For example, if one considers only degrees l = 3-8, the significance levels of the correlations associated with the GTR1 and Mauk (1977) ocean-continent functions reduce to about 55% and 60%, respectively, and hence indicate insignificant correlations. Our conclusion also differs substantially from that of Souriau and Souriau (1983) who, using a Monte Carlo scheme based on random rotations of the data sphere with respect to the tectonic sphere, found that the non-hydrostatic geoid (l = 3-12) correlates significantly (at the 95% confidence level) with the surface tectonics defined by Okal (1977). The close geoid-tectonic association obtained by Souriau and Souriau (1983) is partially related to the fact that the regionalization of Okal (1977) includes subduction zones; the association between the geoid and this regionalization is a result of the strong geoid-slab correlation (e.g., Hager, 1984). Unlike our study, Souriau and Souriau (1983)perform their projections in the spatial rather than in the spherical harmonic domain. After reproducing their results, we repeated
Fig. 6. (a) Histogram of correlations (r) between an ocean-continent function derived from GTR1 and 10,000 synthetic degree-eight fields each with the same spectral properties as the non-hydrostatic geoid. The shaded and unshaded arrows indicate the variance reductions associated with the actual geoid and the maximumvariance reduction, respectively. (b) X2, calculated from the fit of Student's t distribution with the set of t's calculated from t= r[(v/(1- r2)]1/2, plotted as a function of the number of degrees of freedom (v). The minimum value of X2 corresponds with v = 30. Histogram of values of t with Student's t distribution with v degrees of freedom superposed: (c) v = 25, (d) v = 30, and (e) v = 35.
S.S. Shapiro et al.,/ Lithos 48 (1999) 135-152
EI
9~ 1 0 0 50
-0.5
v = 25
0 -5
30
0 -5
149
S.S. Shapiro et al. ,/Lithos 48 (1999) 135-152
150
their suite of projections in the spherical-harmonic domain. We found that the correlation between the long-wavelength geoid (l = 3-12) and the regionalization of Okal (1977) is significant at about the 98% confidence level, slightly higher than the result of Souriau and Souriau (1983) of about 95% from a spatial-domain analysis. However, when we substitute a slab-residual model geoid (Hager and Clayton, 1989) for the geoid, we find that the confidence level reduces to about 50%, indicating that the signal observed by Souriau and Souriau (1983) is largely due to the correlation between slabs and the regionalization. The isopycnic hypothesis (Jordan, 1988) predicts a value of zero for Oln p / O l n v s. This value is within 2.0o- of our estimate and indicates that at this level of significance, the isopycnic hypothesis is consistent with the average geoid anomaly associated with platforms and shields. We can also estimate the value of 01np/Olnu~ by considering only thermal effects on density: Olnp
(1/p)(~p/~T)
Olnus
(l/vs)(Svs/ST)
(6)
Using a coefficient of volume expansion of 3 • 10-5 K - ~, we make two estimates" (1) Oln p/Oln us ~ 0.23, using ~ v s / ~ T ~ - 0.6 m s- ~ K - ~ from McNutt and Judge (1990) and an average upper-mantle shear velocity of ~'s--4.5 km s -1, and (2) Oln p/Olnu s 0.27, using (Olnvs/OT) ~ - 1.1 • 10 - 4 K-1 from Nataf and Ricard (1996). The average of these estimates is inconsistent at about the 10o- level with the value of Oln p/Oln us that we estimate for the continental tectosphere. Hence, our analysis indicates that a simple conversion of shear-wave velocity to density via temperature dependence is inappropriate for the continental tectosphere and that one must consider compositional effects. Our conclusion could not differ more completely from that of Pari and Peltier (1996) (henceforth PP), who claim that they can rule out the hypothesis that neutrally buoyant, compositionally distinct material exists beneath "cratons." Based on a match to the peak amplitude of a severely truncated ( / = 2-8) free-air gravity anomaly at one location (Hudson Bay), they argue that 0.21 < (Oln p/Oln vS) < 0.26, consistent with the thermal estimate above, and in-
consistent at the 8-10o- level with our estimate. However, there are several easily identifiable differences between their approach and ours. Most importantly, we use a geologic regionalization to define cratons. PP define "cratons" as any region, beneath either continents or margins, that has high inferred densities at 30 km depth in heterogeneity model S.F1.K/WM13 (Forte et al., 1994). This definition of "craton" is inappropriate for testing the composition of tectosphere for many reasons, including: (1) 30 km depth beneath continents is generally within the lower crust, not within the proposed isopycnic region of the continental tectosphere; (2) model S.F1.K/WM13 is a heterogeneity model based on a weighted fit both to the gravity field and to the seismic data, assuming that density and velocity anomalies are proportional through assumed depthdependent values of ( O l n p / O l n v s) which vary between 0.21 and 0.34 (Karato, 1993). In regions where the seismic coverage is not good, this assumption introduces a strong gravitational bias into model S.F1.K/WM13, making the use of this model in the inversion for (Olnp/Olnu s) an example of circular logic; (3) The use of this hybrid model fails to identify the South African craton, a region with thick tectosphere (e.g., Su et al., 1994), as a craton. (4) PP emphasize the value of the fit at Hudson Bay, while our study weights all regions of the globe equally. We also note that the estimate of the amount of this peak free-air gravity anomaly attributable to mantle structure is suspect due to contamination from postglacial rebound (Simons and Hager, 1997). In summary, given their approach, and their non-geologic definition of cratons, it is not surprising that PP find a different value for (Oln p/Oln vs) than we do. Their value applies to the mantle beneath regions of inferred high-density lower-crust in a model determined from a joint inversion of gravity and seismic data. Our value of ( O l n p / O l n v s) applies to cratons defined by geological processes. In summary, to obtain realistic estimates of the significance of correlations between data fields defined on a sphere requires that one consider the spectra of the data fields so that the number of degrees of freedom can be determined appropriately. Our analysis demonstrates that the relationship between the long-wavelength geoid and the ocean-continent function is tenuous. The large difference in
s.s. Shapiro et al./Lithos 48 (1999) 135-152
correlation that we obtain with different o c e a n - c o n tinent functions further illustrates its insignificance. F r o m error estimates that account for the redness in the geoid, gravity field, and tectonic regionalization spectra, we conclude that neither the geoid nor the free-air gravity has a platform and shield signal that differs significantly (20-) from zero. Additionally (see Shapiro, 1995), by considering regionally averaged shear-wave travel-time anomalies together with our m o d e l of the continental t e c t o s p h e r e ' s contribution to the geoid, we find that a l n p / a l n v S is about 0.04 _+ 0.02. A l t h o u g h this estimate is consistent at the 2.00- level with the isopycnic hypothesis of Jordan (1988), the slightly positive estimate suggests that the decreased density associated with compositional b u o y a n c y does not completely balance the increased density associated with low temperatures. W e also note that convection calculations addressing the stability and d y n a m i c s of the continental tectosphere indicate that a l n p / a l n u s is likely to vary s o m e w h a t with depth (Shapiro et al., 1999; see also Forte et al., 1995). Thus, our estimate is a weighted average of a quantity that m a y vary with position.
Acknowledgements W e thank P. Puster and G. Masters for c o m p u t e r code and T.A. Herring, P. Puster, W.L. Rodi, and M. Simons for helpful discussions. Richard J. O ' C o n n e l l provided a useful review. Figs. 1, 2 and 5 were created using the Generic M a p p i n g Tools software (Wessel and Smith, 1991). W e p e r f o r m e d m a n y of the calculations using the Guilford College Scientific C o m p u t a t i o n and Visualization Facility which was created with funds from a grant from the National Science F o u n d a t i o n ( C D A - 9 6 0 1 6 0 3 ) . This work was also supported by National Science F o u n d a t i o n grant EAR-9506427.
References Chase, C.G., 1979. Subduction, the geoid, and lower mantle convection. Nature 282, 464-468. Cramer, H., 1946. Mathematical methods of statistics. Princeton Univ. Press, Princeton. Crough, S.T., Jurdy, D.M., 1980. Subducted lithosphere, hotspots and the geoid. Earth Planet. Sci. Lett. 48, 15-22. Doin, M.-P., Fleitout, L., McKenzie, D., 1996. Geoid anomalies
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and the structure of continental and oceanic lithospheres. J. Geophys. Res. 107, 16135-16199. Dziewonski, A.M., Anderson, D.L., 1981. Preliminary reference earth model (PREM). Phys. Earth Planet. Inter. 25, 297-356. Eckhardt, D.H., 1984. Correlations between global features of terrestrial fields. Math. Geol. 16, 155-171. Forte, A.M., Dziewonski, A.M., Woodward, R.L., 1993a. Aspherical structure of the mantle, tectonic plate motions, nonhydrostatic geoid and topography of the core-mantle boundary. In: Le MouEl, J.-L., Smylie, D.E., Herring, T. (Eds.), Dynamics of the Earth's Deep Interior and Earth Rotation. Geophys. Monogr. Ser. 72. AGU, Washington, DC, 135-166. Forte, A.M., Peltier, W.R., Dziewonski, A.M., Woodward, R.L., 1993b. Dynamic surface topography: a new interpretation based upon mantle flow models derived from seismic tomography. Geophys. Res. Lett. 16, 225-228. Forte, A.M., Woodward, R.L., Dziewonski, A.M., 1994. Joint inversions of seismic and geodynamic data for models of three-dimensional mantle heterogeneity. J. Geophys. Res. 99, 21857-21877. Forte, A.M., Dziewonski, A.M., O'Connell, R.J., 1995. Continent-ocean chemical heterogeneity in the mantle based on seismic tomography. Science 268, 386-388. Gurnis, M., 1993. Comment on "Dynamic surface topography: a new interpretation based upon mantle flow models derived from seismic tomography" by A.M. Forte, W.R. Peltier, A.M. Dziewonski and R.L. Woodward. Geophys. Res. Lett. 20, 1663-1664. Hager, B.H., 1983. Global isostatic geoid anomalies for plate and boundary layer models of the lithosphere. Earth Planet. Sci. Lett. 63, 97-109. Hager, B.H., 1984. Subducted slabs and the geoid: constraint on mantle rheology and flow. J. Geophys. Res. 89, 6003-6015. Hager, B.H., Clayton, R.W., Richards, M.A., Comer, R.P., Dziewonski, A.M., 1985. Lower mantle heterogeneity, dynamic topography and the geoid. Nature 113, 541-545. Hager, B.H., Clayton, R.W., 1989. Constraints on the structure of mantle convection using seismic observations, flow models, and the geoid. In: Peltier, W.R. (Ed.), Mantle Convection, Plate Tectonics and Global Dynamics. Gordon & Breach, New York, 657-753. Hager, B.H., Richards, M.A., 1989. Long-wavelength variations in Earth's geoid: physical models and dynamical implications. Philos. Trans. R. Soc. London, Ser. A 328, 309-327. Haxby, W.F., Turcotte, D.L., 1978. On isostatic geoid anomalies. J. Geophys. Res. 83, 5473-5478. Jordan, T.H., 1975. The continental tectosphere. Rev. Geophys. Space Phys. 13, 1-12. Jordan, T.H., 1978. Composition and development of the continental tectosphere. Nature 274, 544-548. Jordan, T.H., 1981. Global tectonic regionalization for seismological data analysis. Bull. Seismol. Soc. Am. 71, 1131-1141. Jordan, T.H., 1988. Structure and formation of the continental tectosphere. J. Petrology, Special Lithosphere Issue, 11-37. Kaula, W.M., 1967. Geophysical applications of satellite determinations of the earth's gravitational field. Space Sci. Rev. 7, 769-794.
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Kaula, W.M., 1972. Global gravity and tectonics. In: Robinson, E.C. (Ed.), The Nature of the Solid Earth. McGraw-Hill, New York, 386-405. Karato, S.-I., 1993. Importance of anelasticity in the interpretation of seismic tomography. Geophys. Res. Lett. 20, 1623-1626. Lemoine, F.G. et al., 1996. The development of the NASA GSFC and NIMA joint geopotential model. Proceedings of the International Symposium on Gravity, Geoid, and Marine Geodesy. Mauk, F.J., 1977. A tectonic based Rayleigh wave group velocity model for prediction of dispersion character through ocean basins. PhD thesis, Univ. of Michigan, Ann Arbor. McNutt, M.K., Judge, A.V., 1990. The superswell and mantle dynamics beneath the south Pacific. Science 248, 969-975. Nakiboglu, S.M., 1982. Hydrostatic theory of the earth and its mechanical implications. Phys. Earth Planet. Inter. 28, 302311. Nataf, H.-C., Ricard, Y., 1996. 3SMAC: an a priori tomographic model of the upper mantle based on geophysical modeling. Phys. Earth Planet. Inter. 95, 101-222. O'Connell, R.J., 1971. Pleistocene glaciation and the viscosity of the lower mantle. Geophys. J. R. Astron. Soc. 23, 299-327. Okal, E.A., 1977. The effect of intrinsic oceanic upper-mantle heterogeneity on regionalization of long-period Rayleigh-wave phase velocities. Geophys. J. R. Astron. Soc. 49, 357-370. Panasyuk, S.V., 1998. The effect of compressibility, phase transformations, and assumed density structure on mantle viscosity inferred from Earth's gravity field. PhD Thesis, MIT. Pail, G., Peltier, W.R., 1996. The free-air gravity constraint on subcontinental mantle dynamics. J. Geophys. Res. 101, 28105-28132. Parsons, B., Richter, F.M., 1980. A relation between the driving
force and geoid anomaly associated with mid-ocean ridges. Earth Planet. Sci. Lett. 51,445-450. Richards, M.A., Engebretson, D., 1992. Large-scale mantle convection and the history of subduction. Nature 355, 437-440. Richards, M.A., Hager, B.H., 1988. The earth's geoid and the large-scale structure of mantle convection. In: Runcorn, S.K. (Ed.), The Physics of Planets. Wiley, 247-272. Richards, M.A., Hager, B.H., 1989. Effects of lateral viscosity variations on long-wavelength geoid anomalies and topography. J. Geophys. Res. 94, 10299-10313. Sclater, J.G., Jaupart, C., Galson, D., 1980. The heat flow through oceanic and continental crust and the heat loss of the Earth. Rev. Geophys. 18, 269-311. Shapiro, S.S., 1995. The stability and dynamics of the continental tectosphere. PhD Thesis, MIT. Shapiro, S.S., Hager, B.H., Jordan, T.H., 1999. Stability and dynamics of the continental tectosphere. This volume. Simons, M., Hager, B.H., 1997. Localization of the gravity field and the signature of glacial rebound. Nature 390, 500-504. Sipkin, S.A., Jordan, T.H., 1975. Lateral heterogeneity of the upper mantle determined from the travel times of ScS. J. Geophys. Res. 80, 1474-1484. Souriau, M., Souriau, A., 1983. Global tectonics and the geoid. Phys. Earth Planet. Inter. 33, 126-136. Su, W.-J., Woodward, R.L., Dziewonski, A.M., 1994. Degree 12 model of shear velocity heterogeneity in the mantle. J. Geophys. Res. 99, 6945-6980. Turcotte, D.L., McAdoo, D.C., 1979. Geoid anomalies and the thickness of the lithosphere. J. Geophys. Res. 84, 2381-2387. Wessel, P., Smith, W.H.F., 1991. Free software helps map and display data. EOS Trans. AGU 72 (441), 445-446.
LITHOS
ELSEVIER
Lithos 48 (1999) 153-170
The evolution of continental roots in numerical thermo-chemical mantle convection models including differentiation by partial melting J.H. de Smet *, A.P. van den Berg ~, N.J. Vlaar
2
Department of Theoretical Geophysics, University of Utrecht, PO Box 80.021, 3508 TA Utrecht, The Netherlands
Received 30 April 1998; received in revised form 21 January 1999; accepted 25 January 1999
Abstract
Incorporating upper mantle differentiation through decompression melting in a numerical mantle convection model, we demonstrate that a compositionally distinct root consisting of depleted peridotite can grow and remain stable during a long period of secular cooling. Our modeling results show that in a hot convecting mantle partial melting will produce a compositional layering in a relatively short time of about 50 Ma. Due to secular cooling mantle differentiation finally stops before 1 Ga. The resulting continental root remains stable on a billion year time scale due to the combined effects of its intrinsically lower density and temperature-dependent rheology. Two different parameterizations of the melting phase-diagram are used in the models. The results indicate that during the Archaean melting occurred on a significant scale in the deep regions of the upper mantle, at pressures in excess of 15 GPa. The compositional depths of continental roots extend to 400 km depending on the potential temperature and the type of phase-diagram parameterization used in the model. The results reveal a strong correlation between lateral variations of temperature and the thickness of the continental root. This shows that cold regions in cratons are stabilized by a thick depleted root. 9 1999 Elsevier Science B.V. All rights reserved. Keywords: Thermo-chemical convection; Numerical modeling; Upper mantle differentiation; Continental evolution; Partial melting; Continental root
1. I n t r o d u c t i o n
Continental nuclei are much older than oceanic lithosphere. Continents have cratonic segments with ages of 3.0 billion years and older (Condie, 1984), in contrast with the oceanic lithosphere with ages up to about 200 million years (Miiller et al., 1996). Thus,
* Corresponding author. E-mail: [email protected]; fax: + 31-30253-50-30 1 E-mail: [email protected]. 2 E-mail: [email protected].
continents apparently form stable systems in the sense that they do not subduct. Oceanic lithosphere is subducting at ocean-continent boundaries, and this process is relatively well understood. The evolution of continental systems is much less clear. In this paper, we present a model for the formation and long term evolution of continental systems within the framework of a numerical mantle convection model. There exists evidence for a specific continental configuration that extends to several hundreds of kilometers into the upper mantle. It has been suggested that the continental lithosphere, also called the
0024-4937/99/$- see front matter 9 1999 Elsevier Science B.V. All fights reserved. PII: S0024-4937(99)00028-6
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J.H. de Smet et al./Lithos 48 (1999) 153-170
tectosphere (Jordan, 1975), is a cold layer that is prevented from collapsing through a compositionally determined low intrinsic density. An overview of evidence supporting such a chemically distinct layer is given by Jordan (1988). Estimates of the depths of these chemically defined continental roots range from 200-400 km (Jordan, 1975, 1988; Anderson, 1990; Polet and Anderson, 1995; Doin et al., 1996). During the Archaean era the Earth had a hotter upper mantle because radiogenic heat production was higher and more initial heat from planetary formation and early differentiation was still stored in the Earth. It has been demonstrated that a hotter geotherm has a large effect on the depth where pressure release partial melting starts in the upper mantle (McKenzie, 1984). This melting process results in residual material with an intrinsically lower density and a basaltic crust derived from the primary melt. This low density residual material is depleted peridotite or harzburgite and its presence can explain the long term stability of ancient, i.e., > 3 Ga, continental areas through gravitationally stable compositional layering. The importance of a higher mantle temperature to the stability of oceanic and continental lithospheric systems has been discussed in Vlaar and Van Den Berg (1991), Vlaar et al. (1994) who used simple models based on the 1-D adiabatic model for partial melting by McKenzie (1984). The results of these 1-D models showed that the present day style of plate tectonics cannot be extrapolated to the Archaean, since the compositionally differentiated layers produced by a convecting mantle must have been much thicker during the Archaean. It was shown that a thick layer of harzburgitic residue underneath an also thicker crust must have led to a different thermo-chemical convection regime (Vlaar et al., 1994). In this paper, we extend these earlier models of partial melting to non-adiabatic conditions within a thermo-chemical mantle convection model (Dupeyrat et al., 1995; De Smet et al., 1998a). We use a fully dynamical numerical mantle convection model including partial melting phase-diagrams based on empirical data. A similar model has been used by De Smet et al. (1998a) and it has been further extended to investigate continental evolution. A new more realistic parameterization of the melting phase-diagram is incorporated in the present work
allowing deep melting for a realistic young Earth continental geotherm. Here, we apply this dynamic mantle differentiation model to investigate an upper mantle system subject to secular cooling in a continental setting and study in particular the formation and subsequent thermo-chemical evolution of the continental root. Other workers have investigated continental systems in the context of mantle convection models. Effects of rheology have been studied by Moresi and Solomatov (1997) and Schmeling and Bussod (1996). Effects of composition and rheology on delamination of a lithospheric root are studied by Schott and Schmeling (1998). Several authors use ad hoc compositional layering of the continental upper mantle. Lenardic (1997) uses this type of model to explain surface heat flow data. Stability aspects are examined by Doin et al. (1997). Long-term evolutions for the whole upper mantle including partial melting are examined by Ogawa (1994) and Kameyama et al. (1996). Walzer and Hendel (1997) included the lower mantle in their study of mantle evolution with respect to the chemical differentiation of heat-producing elements. The Hawaiin hotspot and rifting scenarios in relation with partially melting mantle plumes are studied by Ribe and Christensen (1994), Ebinger and Sleep (1998), and Leitch et al. (1998). In Section 2 we present our model which focuses on the Archaean continental upper mantle. Incorporated are several important physical aspects, such as adiabatic compression, viscous heating, time- and depth-dependent radiogenic heat generation, latent heat consumption, and a pressure- and temperaturedependent rheology.
2. The model
We have applied a numerical convection model for the upper mantle including pressure release melting. A similar model with a different parameterization of the melting phase diagram was applied by De Smet et al. (1998a).
2.1. Conceptual continental model Fig. 1 gives a schematic depth profile of the model. The model incorporates a crust of 30 to 50
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z=0km z= 16km z = 30 or 50 k m
l
u p p e r crust: H c l o w e r crust: H1
d e p l e t e d mantle" H d z-250 km
u n d e p l e t e d mantle: H u
z=670 km Fig. 1. Schematic depth profile of the continental model within the left-hand column depths in km. The upper and lower crust form a high density of viscosity layer with high radiogenic heat sources. The upper part of the mantle consists of a depleted layer up to approximately 250 km depending on the mantle temperature and in the present context also depending on the type of parameterization of the melting phase diagram (see Fig. 2). This depleted layer has an intrinsically lower density ( p(F > 0) < Ps) than the deeper undepleted part of the upper mantle ( p(F = 0) = p,~). km thickness overlaying the upper mantle to a depth of 670 km, which is assumed to be mechanically decoupled from the lower mantle. The upper mantle consists of two layers: a deep lherzolitic layer of undepleted mantle peridotite and a more shallow harzburgitic layer with a variable degree of depletion. The depleted residual material is produced when the lherzolitic source material crosses its solidus and partial melting produces a melt fraction, which is assumed to be extracted instantaneously. The process of partial melting also effects the distribution of radiogenic heat sources H ( Y , t ) . Fig. 1 gives a schematic depth profile of the model
distinguishing four layers with distinct values for H(~',t). The deeper undepleted mantle has an H u = 0.09 p~W m -3, while the partial melting process reduces the value for the depleted layer to H a = 0.04 or 0.0 I~W m -3. This is due to the extraction of heat-producing incompatible elements (Walzer and Hendel, 1997). We used H~ = 0.9 p~W m -3 in the lower and H c = 4.13 ~W m - 3 in the upper crust. The values have been derived from the present day values given by O'Connell and Hager (1980) and Chapman (1986) by applying an amplification factor of two accounting for the mean radiogenic decay since the Archaean. Estimations of Archaean radiogenic heat sources vary from two to three times the present day value. We have adopted the lower bound of this density range and included a decrease of productivity through the radioactive decay constant A = 0 . 3 4 7 Ga -1, which corresponds to a half-life time of 2 Ga (Turcotte and Schubert, 1982). Three different models A to C will be discussed. Model A and B have the heat generation parameterization mentioned above, whereas Model C has zero heat generation when the degree of depletion is larger than zero, i.e., H a = 0 when F > 0 . Both Models A and B have a crustal thickness of 50 km, whereas Model C has a reduced crustal thickness of 30 km. Further differences between the Models A to C are discussed below and an overview of the model characteristics is given in Table 1. The high crustal heat generation values result in a significant crustal blanketing effect. Since the model crust also has a low density ( Pc = 3000 kg m - 3 ) and a high viscosity (Tlmax--- 10 24 Pa s) it is stably situated on top of the continental mantle. Given the constant temperature at the surface (T~rf), the shallow geotherm in the Mechanical Boundary Layer (MBL)
Table 1 Differences between models A, B and C. The last column refers to the two initial geotherms given in Fig. 2a which are both warm and cold is meant in a relative way. The heat generation density for depleted material is given in column four (Ha). All model differences are discussed in Section 2 Model
Crustal thickness (km)
Thermal coupling upper/ lower mantle
Ha (mW m- 3)
Phase-diagram parameterization
Initial geotherm
A B C
50 50 30
reservoir isolated isolated
0.04 0.04 0.0
linear polynomial polynomial
'hot' 'hot' 'cold'
J.H. de Smet et al./Lithos 48 (1999) 153-170
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at the top of the model is to a large extend determined by its heat source distributions. In the deeper region, the geotherm is influenced by several other factors. One of them is the compositional layering. The formation of a depleted low density layer impedes large scale convection and the formation of an adiabatic mantle geotherm at this depth level. An additional factor is the rheological parameterization including a strong temperature dependence of the viscosity. The effect of cooling from the top is a strong increase of the viscosity which
gradually reduces the vigor of convection in the depleted zone which becomes stagnant from the top down. We have used a Newtonian temperature- and pressure-dependent rheology. The viscosity model follows an Arrhenius relation which, in its dimensional form, is:
~1( p,T) = ~ e x p
E+pV RT
(1)
where values are given in Table 2 and the viscosity
Table 2 Explanation of used symbols Symbols
Definition
Value
a P0 p Pc p 77 r/0
Thermal expansion coefficient Reference density Effective density p(T, F ) Crustal density at T = Tsrf Density drop upon full differentiation Non-dimensional viscosity Reference viscosity Maximum viscosity Activation energy Activation volume Gas constant Viscosity pre-factor Degree of depletion
3 • 10- 5 3416 3000 226 1021
T}max
E V R F Y T q ( X ) hor
AT Tsrf
TO T~(p) Tl(P) AT( P)sl p Ap h A K k Cp AS H H0 qb ~' Di Rb Ra
Velocity Non-dimensional temperature Heat-flow Horizontally averaged quantity X Temperature scale Surface temperature Non-dimensional surface temperature T s r f / A T Pressure-dependent solidus temperature Pressure-dependent liquidus temperature Distance between solidus and liquidus Pressure Non-dimensional hydrodynamic pressure Depth scale Radioactive decay constant Thermal diffusivity Thermal conductivity Heat capacity at constant pressure Entropy change upon melting Non-dimensional radiogenic heat generation Reference value radiogenic heat generation Viscous dissipation Internal heating number: H 0 ( h 2 ) / ( k A T ) Dissipation number: agh/cp Compositional Rayleigh number ~ p gh 3/ KrlO Thermal Rayleigh number: Po aATgh3/Krlo
10 24
250 11.0 8.3143 9 x 1011 -
Units K-1
kg m - 3 k g m -3 kg m -3 k g m -3 Pa s Pa s kJ m o l - 1 ixm 3 m o l - 1 J mol-
1 K- 1
Pa s ms
-1
Wm -2 2200 273.15 0.12416 670 0.347 10 - 6 3.416 1000 300 5.33 X 10 -6 318.4 0.1970 0.6661 • 10 6 0.6645 X 106
K K K K K Pa km (aa) -1 m 2 s-1 W m - 1 K-1 J k g - 1 K-1 J k g - 1 K-1 Wm -3 Wm -3
J.H. de Smet et al./Lithos 48 (1999) 153-170
is truncated at its maximum value "r/max = 10 24 Pa s which equates the crustal value. The viscosity prefactor ~ ' is determined from the constraint that at z = 400 km and T = 2023.15 K (i.e., 1750~ the viscosity equals the reference viscosity value r/0 = 10 21 Pas. In the present models we apply a stronger pressure dependence of the rheology, expressed in a activation volume of V = 11.0 txm 3 mol-1 compared to a V = 7.5 txm 3 mo1-1 used by De Smet et al. (1998a). Both numbers are in the range of possible values for the upper mantle (Ranalli, 1991; Karato and Wu, 1993). The increase of the pressure dependence increases the viscosity mainly in the deeper parts of the model, whereas at shallower depth the mobility increases and the zone of minimum viscosity occurs at shallower depth (z = 80 km at t = 0 in the present Models A and B). This results in a thinner MBL and an overall decrease of the geotherm of about 175 K. We consider the reduced maximum temperature in the present models to be more representative for the young Earth. The horizontally averaged geotherm for t = 0 for Models A and B used here is depicted as the thick dotted line in Fig. 2a. The thin dotted line in Fig. 2a is the colder horizontally averaged geotherm corresponding to Model C. The difference is caused by the thinner crust in Model C. In mantle peridotites subject to partial melting, the solubility of water is higher in the basaltic melt fraction than in olivine. Furthermore, hydrous olivine has a lower viscosity then anhydrous olivine. Thus, basaltic melt extraction will increase the viscosity of the residual material (Karato, 1990; Karato and Jung, 1998; Hirth and Kohlstedt, 1996). This effect can amount up to two orders of magnitude as estimated by Karato and Jung (1998). The influence of a modest increase in viscosity, i.e., the viscosity change of almost dry olivine to dry olivine, on the model has been investigated by De Smet et al. (1998a). It stabilizes the root further and increases the thermal blanketing effect slightly. The initial geotherm has been computed from the resulting temperature field obtained from a startup convection run. In this startup run, partial melting and the decrease of heat production in time are artificially switched off. After several hundred million years a statistically steady-state sets in, and a single snapshot of the temperature field of this
157
Fig. 2. Parameterizations used in the models. (a) The thick lines are the solidus and liquidus for the linear (dashed) and polynomial (solid) parameterization where S and L indicate the solidus and liquidus for both parameterizations, respectively. The thinner lines correspond with phase equilibrium lines for which a degree of depletion is reached of 30% and 60%. The kink at 15 GPa corresponds to the phase transition that takes place at this depth. The thick dotted line is the initial geotherm at which melting is switched on at t -- 0 for Models A and B. For Model C, this is the thin dotted line. (b) (inset), degree of depletion ( F ) as function of normalized super-solidus temperature 0 for the two models. The linearized curve (dashed) is a simplification of data given by Jacques and Green (1980) and corresponds with the dashed phase-lines of Fig. 2a. The bent curve (solid) is the third-order polynomial fit as determined by McKenzie and Bickle (1988) and is used in the phase-diagram given by the curved phase-lines of Fig. 2a. The dots correspond to F-values of 0, 30, 60 and 100% and for which the phase lines are given in Fig. 2a.
steady-state is used as an initial condition for a subsequent model run including partial melting and decrease of radiogenic heat production with time. We refer to previous work (De Smet et al., 1998a) for a more extensive discussion.
2.2. Parameterization of the partial melting process Mantle differentiation through partial melting is implemented in our model based on a melting
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phase-diagram for peridotite, which gives the equilibrium value of degree of depletion F for given values of p and T. In previous work (De Smet et al., 1998a) we have used a simple parameterization of the phase-diagram, using linear and parallel curves for solidus and liquidus (Takahashi and Kushiro, 1983). Here, we also apply an improved parameterization based on a higher order polynomial fit of empirical data for the solidus and liquidus of mantle peridotite. Solidus and liquidus lines for both parameterizations are shown in Fig. 2a labeled with S and L, both the linear (dashed lines) and the polynomial (solid lines) parameterizations are shown. Up to 15 GPa the curved solidus and liquidus are third-order fits to data from Gasparik (1990) and Takahashi (1990) also used by Vlaar et al. (1994). For pressures in excess of 15 GPa, a second-order fit to the data from Ohtani et al. (1986) has been applied. The sample material used in both references are not identical and therefore a constant shift of about - 7 0 K has been applied to the data points from Ohtani et al. (1986) such that solidus and liquidus are continuous at p - 15 GPa. With /~= (1,p,p2,p3)r the parameterization used for the solidus is:
rs(p) as - ( 1136, 134.2, - 6 . 5 8 1 , 0 . 1054)r 9 p _< 15 GPa as = (1510.76, 46.27 , - 0.8036, 0.0) r 9 p > 15 GPa
(2) where superscript T is the transposed of the vector. For the liquidus we used: r , ( p ) = a~/~ a~ - ( 1762, 57.46, - 3.487 , 0.0769)r 9 p _< 15 GPa d t - (1470.3025,55.53, - 0 . 9 0 8 4 , 0.0)r. p > 15 GPa
(3) From Fig. 2 we conclude that up to 5 GPa and near the solidus (indicated with S in Fig. 2a) both parameterizations are similar. However, for the larger part of the phase-diagram the differences are significant. Deep melting processes for pressure values exceeding 10 GPa, for example, are excluded by the
linear parameterization for realistic geotherms. Also, the polynomial parameterization is combined with a more realistic parameterization of depletion dependency on the normalized super-solidus temperature 0 given by:
f(p,r) =f
zlT~l(p) =/(0)
(4)
where symbol definitions are given in Table 2. We adopted an empirical relation for f ( 0 ) as given by McKenzie and Bickle (1988), which is based on a third order polynomial fit of available empirical data. In Fig. 2b, this relation is represented by the solid curve and the dash-dotted line is the linear relation used by De Smet et al. (1998a), a linear fit derived from data given by Jacques and Green (1980). The dots on the curves correspond to different degrees of depletion (30%, 60%) for which the equilibrium phase lines are drawn in Fig. 2a. The different phase diagrams result in a different distribution of partial melting. In particular, the more realistic curved solidus will result in melt production at greater depths. The differences in the phase diagrams result in different dynamics of the model. This is caused mainly by the corresponding differences in the compositional buoyancy defined in terms of F and in the differences in latent heat consumption during partial melting. The density effecting the buoyancy is given by the linearized equation of state:
p(T,F) = p0[1 - o ~ ( T - Tsrf)
-
-
f6p/po]
(5)
The EOS (Eq. (5)) includes a linear fit of empirical data specifying p(F) from Jordan (1979), where 8p is given in Table 2. In the present work, two sets of modeling results are discussed which differ in the type of parameterization of the melting phase-diagram. Model A uses the linearized parameterization and Model B and C are based on the higher order polynomial parameterization.
2.3. Thermal coupling between upper and lower mantle For secular cooling models of the upper mantle, thermal coupling between upper and lower mantle is important. Here, we consider two limiting cases implemented in the aforementioned Models A to C.
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Model A includes a simple heat reservoir representing the lower mantle which is assumed isothermal T = TR(t) as used by Kameyama et al. (1996). The reservoir does not contain any internal heating. Its volume is twice the volume of the upper mantle and other physical properties equal those of the upper mantle. The heat extracted from the reservoir is computed from the heat flux through the lower boundary of the numerical model, i.e., the upper/lower mantle interface. The reservoir temperature is used as a time-dependent essential boundary condition in Model A. In Models B and C we consider an upper mantle thermally isolated from the lower mantle. With the heat reservoir approximation we establish a more realistic estimate of the effect of a non-zero heat flow from the lower mantle. Models with a zero heat flux condition at the bottom boundary can be considered as end-member cases, which result in a maximum estimate of the cooling rates of secular cooling of the upper mantle system (De Smet et al., 1998a).
2.4. Governing model equations and numerical methods In the following equations, non-dimensional quantities are used unless explicitly stated otherwise. We used a non-dimensionalization scheme described by Van Den Berg et al. (1993). For an infinite-Prandtl-number fluid, the momentum equation with the thermal and compositional Rayleigh numbers Ra and R b respectively, is given by:
V(rI(VY+ VYr)) - V A p = ( R a T + R b F ) ~
(6)
where ~ is the unit vector in the vertical direction aligned with gravity. The definitions of the symbols are listed in Table 2. The transport equation describing partial melting of a volume of mantle material in terms of the degree of depletion F is"
d F(p,T) dt
=
df( 0 ) dt
(7)
where the right-hand-side describes a source distribution of partial melt production following from the phase diagram discussed above. Recurrent melting is
incorporated in this formulation, meaning that a volume of recycled partially melted material experiences further melting when its super-solidus temperature 0 is in excess of any previously reached value. The energy equation used is based on the Extended Boussinesq Approximation (EBA) (Ita and King, 1994): dT dt
D i ( T + T0)uz-- V Z T + ~ H ( z , t ) Di
AS d F
Ra
Cp
+ - - a s - (T+
dt
(8) where the symbols are defined in Table 2. The last three terms of the right-hand-side are: the radioactive heat generation, viscous dissipation, and latent heat consumption. The second term on the left accounts for cooling and heating due to adiabatic (de-)compression. The effect of viscous heating is relatively small since the vigor of convection is low after the depleted layer has formed. The momentum and energy equations are solved with a finite element method (Van Den Berg et al., 1993) and a time-dependent upwind scheme (SUPG) (Hughes and Brooks, 1979; Segal, 1982) is applied to the latter to improve numerical stability in advection dominated regions of the domain. The transport equation for F is solved using a Method of Characteristics (MoC) (Sotin and Parmentier, 1989; Sparks and Parmentier, 1993). A hybrid scheme using both low and high order interpolations for F in combination with a fourth order Runge-Kutta time-integration for the integration part was applied over a structured grid (De Smet et al., 1998b). In order to limit the amount of numerical diffusion over the whole domain, this unequally spaced grid used for the MoC computations of the degree of depletion field F was much denser than the finite element mesh used, and the grid densities were higher than those used by De Smet et al. (1998a). We used a resolution with grid cells ranging from 1121 to 3116 m in the horizontal direction and 921 to 1688 m in the vertical direction. A Predictor-Corrector timestepping scheme is applied to solve the Equations in the following order: Eqs. (6)-(8).
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J.H. de Smet et al./Lithos 48 (1999) 153-170
3. Numerical modeling results We will focus mainly on the differences in the evolution of Models A and B. Results for Model A are qualitatively similar to earlier results described by De Smet et al. (1998a). The increase of the activation volume in our present models results in a colder geotherm due to a thinner MBL. Up to approximately 200 Ma both models A and B show a similar evolution pattern. After this time deep upper mantle differentiation starts to occur in Model B. We first discuss the early stage before the onset of deep melting in Model B. Next we describe the long term evolution of the continental model with special interest in the stability of the system. A brief description of the dynamically created initial condition has been given above.
3.1. Early evolution From Fig. 3a and 3b a global comparison between model A and B can be made. In the left column the compositional field (degree of depletion F) is shown with upper (white) and lower (black) crustal layers on top of the partially depleted layer shown in color. The lateral variation of the temperature field with respect to the horizontally averaged background value, i.e., T - ( T ) h o r , is shown in the right-hand column. All frames also contain instantaneous flow lines indicating clock- and counterclockwise flows as black and white contour lines respectively. The advection velocities in Model A are higher than those in Model B as can be seen from the number of flow lines within each convection cell. This is mainly due to the differences in the accumulated depleted layers resulting from the different melting parameterizations as will be shown below. The small down-wellings of material with low degree of depletion of < 10% (light-blue to white) in Fig. 3.b.2 are due to re-mixing as will be discussed together with the long term evolution. The lateral temperature variations are similar for both models, with exception of the high T-anomaly above the up-welling part of the convecting cells in Fig. 3a2. This large scale two-cell convection pattern is persistent during a long period of the model
evolution. Fig. 3a and 3b illustrate the relatively short time scale of < 100 Ma in which a compositionally distinct continental root builds up in a convecting mantle after the onset of melting at t = 0. This shows that mantle differentiation is a powerful process in a planetary mantle which is hot enough for the mantle adiabat to intersect the solidus. The rapid build-up of a stably layered system results in a reduced heat transport efficiency and a gradual warming of the deeper half of the model. Fig. 4 shows the depth distribution of several horizontally averaged quantities for four time values after the onset of melting for both Model A (Fig. 4a) and B (Fig. 4b). Frames numbered 1 through 4 correspond to (1) degree of depletion F, (2) temperature T, (3) viscosity ~q(p,T), and (4) root-meansquared velocity Vrms. The averaged profiles for F (Fig. 4al and 4bl) indicate that both models initially evolve in a similar way. The continental root grows due to the mechanism of intermittent small scale diapiric melting events as investigated by De Smet et al. (1998a,b) where the parameterization of Model A was used. The same mechanism is operative in Model B. The depleted layer in model B (Fig. 3bFig. 4b l) is thicker than in Model A where melting occurs in a greater depth interval. This is clearly shown by the difference in pressure where the initial geotherm intersects the solidus lines in Fig. 2a. This pressure difference is about 1 GPa and results after 100 Ma in a continental root extending to approximately 180 and 220 km for model A and B, respectively. Note also the finite values of F in the 350 to 500 km depth range after 200 Ma in Model B (Fig. 4bl). Due to the different F(0)-parameterizations, the transition from the depleted root to the deeper undepleted mantle is sharper, i.e., the slope of (F)hor is steeper for Model B (Fig. 4b 1) than in Model A (Fig. 4.a.1). Note that the maxima of (F)hor are in both cases practically identical. This is explained by the fact that for a 0 of about 0.25, an F-value of about 30% is obtained in both F(O) parameterizations. This is indicated by the proximity of 0 values of the two dots for F = 30% in Fig. 2b. Fig. 4a2, 4b2 show the horizontally averaged geotherms. Both models show an initial increase of temperature with time in the deeper half of the model. This is a result of the rapid build up of a
161
J.H. de Smet et al./Lithos 48 (1999) 153-170
I
. . . .
I
500
. . . .
I
1000
'
'
~
"
i
'
" '~' '"
1500
Fig. 3. The depletion field and lateral variations of the temperature field (i.e., g T = T - { T ) h o r ) at 100 Ma and 200 Ma for both parameterizations of the phase-diagram. Figures labeled a and b correspond to Models A and B, respectively. Black and white contour lines indicate clock and counter clockwise flows, respectively. White F-field contouring corresponds to 9% < F < 10%. The depletion for the polynomial case is larger than in the linear case due to the different F-dependency of 0 (see the inset Fig. 2b). Also the thickness of the depleted layer is slightly thicker when the curved solidus/liquidus are used.
shallow layering inhibiting whole layer upper mantle circulation. This layering consists of cold, and therefore strong material which is part of the MBL. This cold and depleted layer is gravitationally stable with
respect to the deeper parts of the model. A superadiabatic regime develops quickly in this stagnant top layer, indicating the predominance of conductive heat transport over advection. In Model B, the
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J.H. de Smet et al./Lithos 48 (1999) 153-170
~q
li
Fig. 4. For all cases shown in Figs. 3 and 5 horizontally averaged profiles are given: (1) the degree of depletion, (2) the temperature, (3) the viscosity, (4) and the velocity root mean square. The dashed curves in (a2) and (b2) are the solidus and liquidus. Note the numerical instabilities in the ( F)-profile for t = 700 Ma in Model B (b 1), which is caused by the too coarse grid over which the transport equation for F is solved. These wiggles are also expressed by the wash-board effect in Fig. 5b2.
geotherm intersects the solidus at large depths at t ~ 200 Ma and melting is initiated in the lower half of the domain. The latent heat effects of the endothermic phase transition from olivine to spinel at 15 GPa may prevent the geotherm from crossing the solidus above 15 GPa. However, solid-state phase transitions are not included in the present models. After several hundred million years ( ~ 300 Ma) the temperature at the bottom of the upper mantle starts to decrease. These results illustrate the thermal blanketing effect of layered continental systems (Gurnis, 1988), which delay secular cooling. This decrease of the bottom temperature is stronger in Model B, because of the absence of heat influx from the lower mantle, and because latent heat consumption and surface heat flow are higher in Model B.
In Fig. 4a3,b3 viscosity profiles are shown, the crustal viscosity is set by the truncation value ~max ---~ 10 24 Pa s applied for numerical reasons. Effects of initial warming at large depth and cooling from the top are reflected in the temperature-dependent viscosity. Model A develops a viscosity minimum which is slightly more pronounced and at shallower depth than Model B. In Fig. 4a4,b4 we show the Vrm s distribution based on horizontal averaging. Model B shows a strongly reduced vigor of convection in the top half of the model. This is explained by the fact that the low viscosity zone in Model A is effectively positioned below the depleted layer resulting in a relatively mobile top layer of the undepleted mantle. In Model B the low viscosity zone occupies a depth
J.H. de Smet et al. ,/Lithos 48 (1999) 153-170
range with finite value of the degree of depletion F, reducing the vigor of convection.
3.2. Long-term evolution Fig. 4al,bl show that models A and B have a different long-term evolution. In Model A the thickness and volume of the continental root gradually increases with time due to the melting events in the upward flow of the two cell convection pattern which persists in the deep undepleted zone. At the sites of the down-wellings depleted material is slowly dragged into the lower half of the model. A low degree of depletion of about 2% volume average is found in this lower half. In Model B much further depleted material is found in this deep region. Only a small part of it originates from re-mixed depleted material. Most of it is formed at depth since the geotherm has reached and crossed the solidus after about 200 Ma of evolution as shown in Fig. 4.b.2. This results in modest but steady partial melting during the time-span from t = 200 Ma to 850 Ma. The wiggles in F shown in Fig. 4bl below 450 km depth are due to the limited grid resolution in the lower part of the model. Fig. 4a2,b2 show the evolution of the geotherm for Models A and B, respectively. In Model B, the geotherm is stabilized close to solidus at large depths (i.e., p > 15 GPa) during the initial warming stage. This is the result of the occurrence of deep partial melting and the associated latent heat consumption at these depths. In Model A, the averaged temperature at 670 km is steadily increasing and is not bounded by partial melting. Due to shallower melting events the root is growing steadily and conductive heat transport slowly becomes more dominant than advection. The geotherms in Fig. 4a2,b2 also show when the heat reservoir approach is used, i.e., Model A, the inflow of heat from the lower mantle is very small. We therefore conclude that the differences in the results of Models A and B are mainly due to the difference in the phase-diagrams and to a lesser extent to the different thermal coupling with the lower mantle. Fig. 4a3,b3 show the corresponding viscosity profiles during the evolution. In Model B the viscosity at large depths reaches an almost stationary value after 200 Ma since the geotherm be-
163
comes almost stationary. As secular cooling proceeds the MBL at the top grows and the minimum viscosity value increases and its position slowly shifts from 200 to 300 km depth during the 200 to 700 Ma time-span. Model A has a less thick depleted top layer and advection rates are higher than in Model B. This results in a hotter geotherm in the root and a smaller viscosity minimum. These differences in geotherm evolution combined with a different evolution of the buoyant zone result in different convection velocity profiles shown in Fig. 4a4,b4. Fig. 4a4 shows that the vigor of convection in Model A has increased near the bottom of the depleted root. This indicates active melting and corresponding production of buoyant residual material. The thickness of the continental root grows accordingly (Fig. 4al). After 200 Ma velocities decrease as an effect of cooling from the top and a corresponding increase of the viscosity, shown in (Fig. 4a3). Fig. 4b4 shows a downward shift of the velocity maximum and an increase of velocity at greater depth which coincides with the onset of deep melting around 200 Ma. If we define that the transition from root to underlying mantle is situated at approximately F = 0.1, we see that at 700 Ma the continental root in Model B has grown to about 400 km thickness with large lateral variations (Fig. 5). Convection velocities start to increase again while the melt production in the deep layer ( p > 15 GPa) continues. Calculations for Model B were stopped at 850 Ma because the insufficient resolution in the bottom layer and corresponding oscillations in the F-field solutions produce increasingly unreliable results beyond that time. Fig. 5a and 5b are contour plots of the depletion and temperature fields for Models A and B, respectively. These snapshots correspond to the profiles as given in Fig. 4, i.e., 400 and 700 Ma. For Model A also the snapshot at t = 1200 Ma is given. In Model A the pattern of convection changes from a relatively vigorous two-cell pattern at 400 Ma to a multi-cell convection regime at 700 Ma, in line with the observed (Vrms >hor profiles discussed above (Fig. 4a4). Melting has stopped at ~ 650 Ma, so the depletion F is subject to advection only from that time on. Both snapshots of Model A show structures of depleted material with depletion values up to 10%
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J.H. de Smet et al./Lithos 48 (1999) 153-170
j
n
1000
]
I
I
'
!
!
,I
I
Fig. 5. As in Fig. 3 but now for t = 400, 700 Ma and for Model A also at 1200 Ma. Figures labeled (a) and (b) correspond to Models A and B, respectively. Due to melting at large depths in Model B the lower half of the domain is also depleted. The wash-board effects visible in the degree of depletion of Fig. 3b2 are numerical instabilities caused by the MoC-method applied over the too coarse grid in this lower part of the domain.
J.H. de Smet et al./Lithos 48 (1999) 153-170
(white), which are dragged down into the undepleted zone. This process contaminates the deeper regions with depleted material at very slow rates. At t = 1200 Ma (Fig. 5a3) there is a small thread-like structure of depleted material (green, i.e., F ~ 15%) delaminating at x = 1000 km. Fig. 5a3 shows more stream-lines at t = 1200 Ma than Fig. 5a2 at t = 700 Ma, indicating that convection rates have increased over this time interval. This is caused by continued cooling form the top, which slowly increases the temperature difference across the deeper undepleted layer and increases convection rates. It also neutralizes the positive compositional buoyancy of the deep part of the continental root, which allows for intermittent small scale delamination (Fig. 5a3). At the same time, however, the MBL extends to larger depths due to the cooling from the top (Fig. 4a3), which prevents the gross part of the root from sinking into the slowly convecting undepleted layer. The lateral temperature variations shown in Fig. 5al,a2 also express the change in convection style. The high T-anomaly above the ascending flow in the two-cell convection decreases from t = 400 to 700 Ma. Fig. 5bl,b2 show the same snapshots for Model B. The scarcity of the streamlines illustrates that the vigor of convection is much lower compared to Model A, which is in line with the 1-D velocity profiles shown in Fig. 4b4. The differences in the depletion fields between Models A and B are large due to the ongoing deep melt production in the latter. This is illustrated by the large amount of depleted material with F > 10% (white = 10%) advected by the convection in the deep zone. Ongoing melt production keeps adding residual material in the continental root which grows to a depth of about 400 km. The 'wash-boarding' effect in the low depletion zone (light blue to white) correspond to the wiggles already described in the 1-D profiles in Fig. 4bl. In Fig. 5a3, no 'wash-boarding' is observed since the used resolution is much higher in the lower part of the domain in Model A than in Model B. The amplitude of the temperature anomalies in Model B increase from t = 400 to 700 Ma as shown in Fig. 5bl and 5b2. This coincides with a temporal acceleration of the convective flow at about t = 700 Ma, related to the up-welling near x = 1000 km and
165
downward flow at x = 500 and 1500 km. The temperature anomalies shown in the fight-hand column correlate with the distribution of composition and the structure of the convective flow. The depleted compositional root is generally cold with the exception of young newly formed residual material in up-welling flow. Both the spatial variations of depletion and of the temperature field are reflected in the geophysical observables: the seismic wave velocity structure of continental areas as revealed by surface wave tomography (Muyzert, 1996; Curtis et al., 1998), wave velocity patterns (Jordan, 1975; Anderson, 1990), and the shear wave velocity distribution on a global scale (Zhang and Tanimoto, 1993), and in the observed gravity field over continental areas (Matyska, 1994; Doin et al., 1996).
3.3. Patterns of mantle differentiation Fig. 6al and 6bl show the evolution of the melt production accumulated over columns in the models as a function of the x-location and the evolution of this melting pattern. The first 50 Ma of the evolution corresponds to the initial phase of rapid differentiation immediately following the onset of melting at t = 0. During this initial period most of the continental root is formed in our model. The figures show that most of the melt is formed in vertical columns with a steady position at the location of warm upwellings in the large-scale convection pattern of the deep mantle shown in Fig. 3. Fig. 6a2,b2 show the evolution accumulated over rows of the computational domain. These frames illustrate the evolution of the melt generation at different depths. Melt production occurs in a wider range for Model B as a consequence of the different shape of the peridotite solidus (Fig. 2a). Deep melt production for p > 15 GPa starts around 200 Ma in Model B (see Fig. 6b) and continues until the end of the computation at 850 Ma. The deep melting has a less pulsating character in time because this type of melting does not take place in rapid vertical ascending flows. Melt generation is concentrated at 600 km depth and near the kink of the phase equilibrium at 448 km depth, i.e., at p = 15 GPa. This is explained by the way in which the almost adiabatic deep geotherm intersects the curved solidus in the this part of the model. The large
166
de Smet et al./Lithos 48 (1999) 153-170
al
I
'''l
....
I ....
I ....
I
....
I~
[ J
,;
-
|
400
-
I
....
t ....
t , . ...... I. . . . . .
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I
b2 100
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200 -
..... '+.... " '
l
._~ 300 c...
~. 4oo "o 500 600
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-
|
....
-
t . . . . [ . . . . I . . . . I,,,I,I
Fig. 6. The horizontal ( a l - c l ) and the 1-D depth ( a 2 - c 2 ) melt production distribution as a function of time for Model A, B and C. Figures labeled (a), (b) and (c) correspond to Models A, B and C, respectively. Note the large difference in partial melting depths for both cases. In (b2) and (c2), the solid line is the depth where the 15 GPa kink in the solidus is located. The dashed vertical lines indicate the time where computations have been stopped. Model calculations for Model A exceed the here given time-window.
melting event between 200-300 km depth at 700 Ma shown in (Fig. 6b2) is reflected in the F- and T-fields shown in (Fig. 5b2). The same melt production plots for Model C are given in Fig. 6cl,c2. Model C has the same phasediagram parameterization as Model B, but no deep melting occurs due to the lower potential temperature and the lower density of heat generating elements. The pattern of melting is therefore similar as in Model A and differentiation takes place at pressures lower than 10 GPa. The depth range is, however, larger than in Model A and extends to a maximum of 250 km. When cooling proceeds, melting ceases as in Model A, although in Model C melting proceeds longer than in Model A. A 3-D extension of the 2-D model will not alter the observed processes to a great extend. On the one
hand, a cylindric type of upwelling in a 3-D model will probably melt further since it can penetrate more easily into the existing root. On the other hand, excess heat contained in the diapir is also lost more rapid in a cylindric configuration, which results in more modest degrees of melting. 3.4. Global evolution
Fig. 7 shows several globally averaged quantities for Models A, B and C, denoted by short dashed, solid, and long dashed lines respectively. The first - 1 5 0 to 0 Ma show the statistically steady-state of the start-up scenario during which partial melting and decay of radioactive heat sources is switched off. Melting is allowed for times larger than t = 0 Ma, when radioactive decay starts. This results in the
J.H. de Smet et al./Lithos 48 (1999) 153-170
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i ....
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....
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C 3.0 ~ ' 2.5
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. . . .
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88 , , , l J , , , l , l l l l l j L ~ [ l l l ~
d
E ......
I .... i''"1 .... I .... I .... i"
A
v
60
,i
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Fig. 7. For both models the evolution of some volume averaged quantities are given. The results for Model A, B and C drawn with short dashed, solid, and long dashed lines, respectively. Volume averages of the following quantities are depicted: (a) temperature, (b) degree of depletion, (c) velocity root-mean-square (Model C not shown). The averaged surface heat flow number is depicted in (d).
short period of about 50 Ma of extensive differentiation illustrated also in Fig. 6, during which a large part of the continental root is formed. The rapid initial formation is an artifact of the particular start scenario used. Subsequent evolution however shows more realistic aspects of the model. The volume averaged temperature shown in Fig. 7a diverges between Model A and B, especially from about 200 Ma onwards. This is mainly due to the effect of latent heat consumption, which increases strongly at the onset of deep melt production in Model B. The secular cooling shown in Fig. 7a is
167
mainly due to conductive cooling from the top, since deeper parts of the model experience an initial warming during several hundred million years as shown in the evolution of vertical profiles in Fig. 4. Model C has an averaged initial temperature which is 70 K lower than in Models A and B. After the onset of melting, owing to the reduced heat production, the temperature in Model C drops much faster than in the two others. Fig. 7b shows the volume average of the degree of depletion. The initial phase of rapid melting is reflected in the steep increase of < F ) . In Model B, more melt is produced and melting continues until the end of the computations whereas in Model A, melting ceases around t = 650 Ma. Model C produces approximately the same volume of continental root as Model A which is explained by the difference in phase-diagram parameterization: in spite of the colder geotherm it can still generate a large amount of depleted material. Fig. 7c shows the root mean squared velocity Vrms based on the volume average. After an initial spike shortly after t - 0 of 5 cm a-1 for Models A and B, truncated in the time series plot, the < VrmS) drops from 2.8 cm a-a during the statistically steady-state before the onset of melting to a value lower than 1 cm a-1. Small local maxima correspond to pulsating diapiric events which also coincide with small shifts in the time series of < F ) (i.e., increasing) and
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J.H. de Smet et al. / Lithos 48 (1999) 153-170
The bottom part of the model forms a relatively slowly convecting layer, and therefore has no pronounced Thermal Boundary Layer (TBL) (Fig. 4a4). Heat flow density through the Earth's surface is shown in Fig. 7d. Values decrease considerably after the formation of a stable continental root that prevents whole layer convection. During later evolution the strong MBL grows from the top down and a more conductive regime, with reduced heat flow, develops. The increase in ( q ) for Model B corresponds to the acceleration in the convective heat flow at approximately t = 700 Ma discussed above. The heat influx from the lower mantle in Model A is less than 7.5% of the heat outflow through the surface, whereas Model B has a zero heatflow contribution from the lower mantle. Nevertheless, the surface heat flow in Model B is slightly higher than in Model A. This is due to the modestly higher advection rates in the very shallow part of the root in Model B compared to Model A (Fig. 4a4,b4), which transports heat more efficiently through this thin shallow layer. Continental heat flow values range between approximately 25 and 75 mW m -2 at present for cratons (Pollack et al., 1993) and the here observed values should reach this range after approximately 2 to 3 Ga which is reasonable. Model C follows Model A but with an approximately constant heat flux difference of 15 to 20 mW m -2. This is mainly due to the reduced radiogenic heat production in the crust of Model C.
4. Concluding remarks Our numerical modeling results show that a hot convecting mantle will produce a compositional layering which is similar to the layered model of the continental tectosphere (Jordan, 1975). It is found that this layering remains stable on a time scale greater than one billion years. These results support the conclusion that present-day cratonic continental roots revealed by seismology are relicts of a hotter Archaean mantle. We have investigated different parameterizations of the melting phase diagram of mantle peridotite. For the more realistic parameterization, we found an increase in melt production which also occurs over a larger depth range, compared to the simpler linear parameterization. This model also shows significant
partial melting in the transition zone of the upper mantle ( p > 15 Gpa), for a sufficiently high initial potential temperature (Model B). This model needs to be explored further in the future. In particular the solid state exothermic phase transition near 400 km depth and the associated latent heat effects on warm diapiric up-wellings and the occurrence of deep melting requires further investigation. Simulations with a simple melting phase diagram show a compositional layering that becomes more pronounced until partial melting stops at about 650 Ma, which result from a progressive secular cooling that eventually brings the average geotherm below the solidus. This resulting layering remains stable for the duration of the computations (1400 Ma), but some small-scale delamination of the continental root and re-mixing of depleted material is observed. In the interpretation of the observed long-term stability of the continental layering, we should keep in mind some limitations of the model. In the real Earth, continents break up. This may be related to the interaction of large scale mantle plumes with continental roots and a certain degree of freedom for continental blocks to migrate laterally (Ebinger and Sleep, 1998). In our model, lateral migration is limited by the symmetry condition applied on the vertical boundaries of the computational domain. At the same time, only relatively small mantle diapirs interact with the continental root and they cannot sufficiently affect the root to result in continental break up. The continental root grows by means of relatively small scale diapiric up-wellings which cross the solidus near the lower boundary of the root. These diapiric flows are then accelerated by the density reduction effect of the partial melting process on the residual matrix, resulting in a positive buoyancy which dominates the effects of latent heat consumption and adiabatic decompression, both included in our models. The amount of melt produced in a single diapiric event is significantly smaller than the estimated amounts produced during events that are related to continental flood basalts. A clear correlation was found between lateral variations of the temperature and of the thickness of the continental root. Cold areas coincide with an increased thickness of the root and warm areas show a thinned root.
J.H. de Smet et al./Lithos 48 (1999) 153-170
Acknowledgements We thank Shijie Zhong and Rob van der Hilst for critical reviews which helped to improve the manuscript. We acknowledge stimulating discussions with David Yuen. This work was partly supported by the Netherlands Organization for Scientific Research (NWO) and the Dr. Schiirmann foundation through travel grants and was partly carried out during a visit of Jeroen de Smet and Arie van den Berg at the Minnesota Supercomputer Institute. Support was also received from the Netherlands Science Foundation (NWO) and NATO.
References Anderson, D.L., 1990. Geophysics of the continental mantle: a historical perspective. In: Menzies, M. (Ed.), Continental Mantle. Oxford, United Kingdom, Clarendon Press, pp. 1-30. Chapman, D.S., 1986. Thermal gradients in the continental crust. In: Dawson, J.B., Carswell, D.A., Hall, J., Wedepohl, K.H. (Eds.), The Nature of the Lower Continental Crust. Vol. 24, Spec. Publ. Geol. Soc. London, pp. 63-70. Condie, K.C., 1984. Archaean geotherms and supra crustal assemblages. Tectonophysics 105, 29-41. Curtis, A., Dost, B., Trampert, J., Snieder, R., 1998. Eurasian fundamental mode surface wave phase velocities and their relationship with tectonic features. J. Geophys. Res. 103, 26919-26947. De Smet, J.H., Van Den Berg, A.P., Vlaar, N.J., 1998a. Stability and growth of continental shields in mantle convection models including recurrent melt production. Tectonophysics 296, 1529. De Smet, J.H., Van Den Berg, A.P., Vlaar, N.J., 1998b. A characteristic based method for solving the transport equation and its application to the process of mantle differentiation and continental root growth (submitted to Geoph. J. Int.). Doin, M.-P., Fleitout, L., Christensen, U., 1997. Mantle convection and stability of depleted and undepleted continental lithosphere. J. Geophys. Res. 102, 2771-2787. Doin, M.-P., Fleitout, L., McKenzie, D., 1996. Geoid anomalies and structure of continental and oceanic lithospheres. J. Geophys. Res. 101, 16119-16135. Dupeyrat, L., Sotin, C., Parmentier, E.M., 1995. Thermal and chemical convection in planetary mantles. J. Geophys. Res. 100, 497-520. Ebinger, C.J., Sleep, N.H., 1998. Cenozoic magmatism throughout East Africa resulting from impact of a single plume. Nature 395, 788-791. Gasparik, T., 1990. Phase relations in the transition zone. J. Geophys. Res. 95 (B10), 15751-15769. Gurnis, M., 1988. Large-scale mantle convection and the aggregation and dispersal of supercontinents. Nature 332, 695-699.
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Hirth, G., Kohlstedt, D.L., 1996. Water in the oceanic upper mantle: implications for rheology, melt extraction and the evolution of the lithosphere. Earth Planet. Sci. Lett. 144, 93-108. Hughes, T.J.R., Brooks, A., 1979. A multidimensional upwind scheme with no crosswind diffusion.In: Hughes, T.J.R. (Ed.), Finite Element Methods for Convection Dominated Flows, Vol. 34 of Applied Mechanics Division, American Society of Mechanical Engineers, New York. Ita, J., King, S.D., 1994. Sensitivity of convection with endothermic phase change to the form of governing equations, boundary conditions and equation of state. J. Geophys. Res. 99, 15919-15938. Jacques, A.L., Green, D.H., 1980. Anhydrous melting of peridotite at 0-15 kb pressure and genises of tholeiitic basalts. Contrb. Mineral. Petrol. 73, 287-310. Jordan, T.H., 1975. The continental tectosphere. Geophys. Space Phys. 133, 1-12. Jordan, T.H., 1979. Mineralogies, densities, and seismic velocities of garnet lherzolites and their geophysical implications. In: Boyd, F.R., H.O.A., Meyer (Eds.), The Mantle Sample: Inclusions in Kimberlites and Other Volcanics, American Geophysical Union, Washington DC, pp. 1-14. Jordan, T.H., 1988. Structure and formation of the continental tectosphere. J. Petrol., Special Lithosphere Issue, 11-37. Kameyama, M., Fujimoto, H., Ogawa, M., 1996. A thermo-chemical regime in the upper mantle in the early Earth inferred from a numerical model of magam-migration in a convecting upper mantle. Phys. Earth Planet. Inter. 94, 187-215. Karato, S., 1990. The role of hydrogen in the electrical conductivity of the upper mantle. Nature 347, 272-273. Karato, S., Jung, H., 1998. Water, partial melting and the origin of the seismic low velocity and high attenuation zone in the upper mantle. Earth Planet. Sci. Lett. 157, 193-207. Karato, S., Wu, P., 1993. Rheology of the upper mantle: a synthesis. Science 260, 771-778. Leitch, A.M., Davies, G.T., Wells, M., 1998. A plume head melting under a rifting margin. Earth and Planetary Science Lett. 161, 161-177. Lenardic, A., 1997. On the heat flow variation from archean cratons to Proterozoic mobile belts. J. Geophys. Res. 102, 709-721. Matyska, C., 1994. Topographic masses and mass heterogeneities in the upper mantle. Gravimetry and Space Techniques Applied to Geodynamics and Ocean Dynamics, Vol. 17 of Geophysical Monograph Series, American Geophysical Union, Washington, DC, pp. 125-132. McKenzie, D., 1984. The generation and compaction of partially molten rock. J. Geophys. 25, 713-765. McKenzie, D., Bickle, M., 1988. The volume and composition of melt generated by extension of the lithosphere. J. Geophys. Res. 29, 625-679. Moresi, L., Solomatov, V., 1997. Mantle convection with a brittle lithosphere: thoughts on the global tectonic styles of the Earth and Venus. Geophys. J. Int., (submitted). Mi~ller, R.D., Roest, W.R., Royer, J.Y., Gahagan, L.M., Sclater, J.G., 1996. Age of the ocean floor. Technical Report MGG-12,
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World Data Center A for Marine Geology and Geophysics, 325 Broadway, Boalder, CO 80303-3328, USA. Muyzert, E., 1996. A seismic cross section through the east European continent. Geophys. J. Int. (submitted). O'Connell, R.J., Hager, B.H., 1980. On the thermal state of the Earth. In: Dziewonski, A.M., Boschi, E. (Eds.), Phys. Earth's Int. Elsevier, New York, pp. 270-317. Ogawa, M., 1994. Effects of chemical fractionation of heat-producing elements on mantle evolution inferred from a numerical model of coupled magmatism-mantle convection system. Phys. Earth Planet. Int. 83, 101-127. Ohtani, E., Kato, T., Sawamoto, H., 1986. Melting of a model chondritic mantle to 20 GPa. Nature 322, 352-353. Polet, J., Anderson, D.L., 1995. Depth extent of cratons as inferred from tomographic studies. Geology 23 (3), 205-208. Pollack, H.N., Hurter, S.J., Johnson, J.R., 1993. Heat flow from the Earth's interior: analysis of the global data set. Rev. Geophys. 31,267-280. Ranalli, G., 1991. The microphysical approach to mantle rheology. In: Sabadini, R. et al. (Eds.), Glacial Isostasy, Sea-Level and Mantle Rheology, Kluwer, The Netherlands, pp. 343-378. Ribe, N.M., Christensen, U.R., 1994. Melt generation by plumes; a study of Hawaiin volcanism. J. Geophys. Res. 99, 669-682. Schmeling, H., Bussod, G.Y., 1996. Variable viscosity convection and partial melting in the continental asthenosphere. J. Geophys. Res. 101,5411-5423. Schott, B., Schmeling, H., 1998. Delamination and detachment of a lithospheric root. Tectonophysics (submitted). Segal, A., 1982. Aspects of numerical methods for elliptic singular perturbation problems. SIAM J. Sci. Stat. Comput. 3 (3), 327-349.
Sotin, C., Parmentier, E.M., 1989. Dynamical consequences of compositional and thermal density stratification beneath spreading centers. Geophys. Res. Lett. 16, 835-838. Sparks, D.W., Parmentier, E.M., 1993. The structure of three-dimensional convection beneath oceanic spreading centers. Geophys. J. Int. 112, 81-91. Takahashi, E., 1990. Speculations on the Archean mantle: missing link between komatiite and depleted garnet peridotite. J. Geophys. Res. 95, 15941-15954. Takahashi, E., Kushiro, I., 1983. Melting of a dry peridotite at high pressure and basalt magma genesis. Am. Miner. 68, 859-879. Turcotte, D.L., Schubert, G., 1982. Geodynamics; Applications of continuum physics to geological problems. Wiley. Van Den Berg, A.P., Van Keken, P.E., Yuen, D.A., 1993. The effects of a composite non-Newtonian and Newtonian rheology on mantle convection. Geophys. J. Int. 115, 62-78. Vlaar, N.J., Van Den Berg, A.P., 1991. Continental evolution and Archeao-sea-level. Sabadini, R. et al. (Eds.), Glacial Isostasy, Sea-Level and Mantle Rheology, Elsevier, The Netherlands, pp. 637-662. Vlaar, N.J., Van Keken, P.E., Van Den Berg, A.P., 1994. Cooling of the Earth in the Archaean; consequences of pressure-release melting in a hotter mantle. Earth Planet. Sci. Lett. 121, 1-18. Walzer, U., Hendel, R., 1997. Time-dependent thermal convection, mantle differentiation and continental-crust growth. Geophys. J. Int. 130, 303-325. Zhang, Y.-S., Tanimoto, T., 1993. High-resolution global upper mantle structure and plate tectonics. J. Geophys. Res. 98, 9793-9823.
LITHOS ELSEVIER
Lithos 48 (1999) 171-194
The age of continental roots D.G. Pearson
*
Department of Geological Sciences, Durham University, South Road, Durham, DH1 3LE, UK Received 20 March 1998; received in revised form 8 January 1999; accepted 11 January 1999
Abstract Determination of the age of the mantle part of continental roots is essential to our understanding of the evolution and stability of continents. Dating the rocks that comprise the mantle root beneath the continents has proven difficult because of their high equilibration temperatures and open-system geochemical behaviour. Much progress has been made in the last 20 years that allows us to see how continental roots have evolved in different areas. The first indication of the antiquity of continental roots beneath cratons came from the enriched Nd and Sr isotopic signatures shown by both peridotite xenoliths and inclusions in diamonds, requiring isolation of cratonic roots from the convecting mantle for billions of years. The enriched Nd and Sr isotopic signatures result from mantle metasomatic events post-dating the depletion events that led to the formation and isolation of the peridotite from convecting mantle. These signatures document a history of melt- and fluid-rock interaction within the lithospheric mantle. In some suites of cratonic rocks, such as eclogites, Nd and Pb isotopes have been able to trace probable formation ages. The R e - O s isotope system is well suited to dating lithospheric peridotites because of the compatible nature of Os and its relative immunity to post-crystallisation disturbance compared with highly incompatible element isotope systems. Os isotopic compositions of lithospheric peridotites are overwhelmingly unradiogenic and indicate long-term evolution in low R e / O s environments, probably as melt residues. Peridotite xenoliths from kimberlites can show some disturbed R e / O s systematics but analyses of representative suites show that beneath cratons the oldest Re depletion model ages are Archean and broadly similar to major crust-forming events. Some locations, such as Premier in southern Africa, and Lashaine in Tanzania, indicate more recent addition of lithospheric material to the craton, in the Proterozoic, or later. Of the cratons studies so far (Kaapvaal, Siberia, Wyoming and Tanzania), all indicate Archean formation of their lithospheric mantle roots. Few localities studied show any clear variation of age with depth of derivation, indicating that > 150 km of lithosphere may have formed relatively rapidly. In circum-cratonic areas where the crustal basement is Proterozoic in age kimberlite-derived xenoliths give Proterozoic model ages, matching the age of the overlying crust. This behaviour shows how the crust and mantle parts of continental lithospheric roots have remained coupled since formation in these areas, for billions of years, despite continental drift. Orogenic massifs show more systematic behaviour of R e - O s isotopes, where correlations between Os isotopic composition and S or Re content yield initial Os isotopic ratios that define Re depletion model ages for the massifs. Ongoing S r - N d - P b - H f - O s isotopic studies of massif peridotites and new kimberlite- and basalt-borne xenolith suites from new areas, will soon enable a global understanding of the age of continental roots and their subsequent evolution. 9 1999 Elsevier Science B.V. All rights reserved.
Keywords: Continent; Lithospheric mantle; Os isotope; Geochronology; Xenolith
* Fax: +44-191-374-2510 0024-4937/99/$ - see front matter 9 1999 Elsevier Science B.V. All rights reserved. PII: S 0 0 2 4 - 4 9 3 7 ( 9 9 ) 0 0 0 2 6 - 2
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D.G. Pearson / Lithos 48 (1999) 171-194
1. W h y date lithospheric mantle?
Much of our understanding of plate tectonics and Earth processes is based on observations from crustal rocks, simply because they are the most immediately accessible but also because the context for interpreting geochemical data for the crust is much clearer. In fact, some of the first positive evidence concerning the antiquity of the continental lithospheric mantle (CLM) came from observations of isotopic diversity in erupted basalts (Brooks et al., 1976). Age constraints on the continental lithosphere address a large variety of questions such as: (i) When were the first stable continental roots formed and did these coincide with the first stable cratonic crust, i.e., does the age of the crust match that of the mantle below? (ii) How was the first CLM formed and has this process varied through time? (iii) Is the formation of crust and CLM linked? (iv) To what extent are the crust and CLM coupled during plate movement and collision? Various models for CLM growth can be envisioned, some of which are illustrated in Fig. 1. Some models might produce distinctive age structure in the CLM that may enable us to discriminate between endmembers (Fig. 1), however, if hybrid models dominate then the picture is probably very difficult to decipher. One of our tasks has been to address this question. Samples of CLM are available for study in two main forms. One is as large fragments of tectonically emplaced peridotites, known as orogenic peridotite massifs (Menzies and Dupuy, 1991b). Such massifs commonly occur at plate-boundaries. The other mode of occurrence is as small fragments of mantle entrained in volcanic rocks during eruption; mantle xenoliths (Nixon, 1987). In contrast to massifs, mantle xenoliths usually sample intra-plate lithospheric mantle and so each sample type has merits over the other, but between them they provide us with a fairly comprehensive sampling of CLM. For instance, whereas massif peridotites form samples of over 100 km 2 in some instances, allowing long-range spatial variability to be studied, xenolith suites, although much more limited in their lateral sampling, may provide fragments from the entire lithospheric depth (Fig. 2).
1 aa 0.5 Ga
~
a~ ~ o-"~_~ ~o
2.4 Ga
._._~ 0.5 Ga
o.5 aa~"~...~...3.~o.5 aa Lateral "block" accretion g~43 0~2 ~
2
Fig. 1. Hypothetical models of continental root formation with possible age relationships. Hybrids between the different endmember models can also be envisaged.
The array of calculated equilibration pressures and temperatures illustrated in Fig. 2 suggest that some lithospheric xenoliths are derived from depths of at least 200 km. In fact, such plots provide a first order view of the potential antiquity of cratonic lithosphere in that the P/T arrays for xenoliths from both Cretaceous and Proterozoic kimberlites (Premier) indicate no change in lithospheric thickness over this time period.
D.G. Pearson/Lithos 48 (1999) 171-194
A
P GPa
173
least the upper portions of the continental crust. Neither of these two criteria are likely to be satisfied in rocks originating from the lithospheric mantle and it is unlikely that the geochronology of the continental (or oceanic) lithospheric mantle will ever be as well constrained and precisely known as that of the overlying crust. This is largely due to one major factor readily evident in Fig. 2: temperature. The high ambient temperature of most of the CLM, under even cool cratonic geothermal conditions, is probably above the blocking temperature of most isotopic systems in most minerals that could be used for multi-mineral or rock isochron-type dating. In addition, these high temperatures promote reaction between mantle and infiltrating fluids such that disturbance of isotopic systems is readily achieved. Also, the scarcity of phases such as zircon in the mantle limits the application of U - P b techniques in accessory phases, which has done so much to resolve problems in crustal geochronology. Despite these problems, some remarkable progress has been made.
-
-
lib
2. Isotopic signatures of C L M
Fig. 2. P~ T plot showing equilibration conditions of southern
African xenoliths compared to the likely blocking temperatures for commonly used isotopic systems. Data from Danchin (1979), Finnerty and Boyd (1987), and Boyd (unpublished). The "kink" in the P / T arrays shown here is not present if some other thermobarometer combinations are employed. See Dickin (1995), and Heaman and Parrish (1991) for blocking temperature estimates.
One of the problems of obtaining age information on samples of CLM using radiogenic isotope systems has been the very low levels of the elements of interest, and so application of these techniques did not really proliferate until the development of modem mass spectrometers in the late 1970s. To understand the genesis and evolution of continental roots it is desirable to have similar stratigraphic and geochronological constraints to those available for at
It is not within the scope of this work to review the basics of the radiogenic isotope systems commonly used in the study of CLM and readers are referred to relevant texts, e.g., Faure (1986) and Dickin (1995). The contribution of S r - N d - P b isotope systems to the characterisation and origin of CLM has been thoroughly reviewed by Menzies (1990). It is worth noting some features of the Re-Os system before proceeding. The advantages of the Re-Os isotope system for obtaining age information were emphasised by Walker et al. (1989a) and Shirey and Walker (1998). The main attractions to this system are discussed below. (i) Os is abundant in mantle rocks relative to either their host magmas, in the case of xenoliths, or most post-crystallisation metasomatic fluids/melts, in the case of both xenoliths and orogenic massifs (Fig. 3), such that it is more robust to disturbance than isotope systems based on incompatible elements (elements concentrated in the liquid phase during partial melting). (ii) Minimum ages of separation of peridotites from the convecting upper mantle can still be ob-
174
D.G. Pearson/Lithos 48 (1999) 171-194
Fig. 3. Illustration of the effects of simple mixing of a peridotite xenolith and host kimberlite magma on the Os isotope composition of the bulk rock xenolith and the Nd isotope composition of diopside from the xenolith (the bulk rock is usually measured for Os isotopes and a diopside or garnet separately measured for the Nd isotope composition). The effects of older interaction, at 1.1 Ga, followed by time-integrated in-growth of radiogenic Os from introduced Re is also modelled. Os isotopic compositions lying to the left of the dashed line give Archaean TRD ages. It is noteworthy that large amounts of kimberlite interaction are required to destroy the signature of Arcaean depletion in a peridotite. See Fig. 13 caption for definition of TRo age.
")/Os(t)= [ (1870S/ 188OSsample(t))/ (1870S/1 880Schondrite(,))] ENd(t ) =
x 100. [ (143Nd/144Ndsample(t))/(143Nd/144 Ndchondrite(t))] • 10000.
tained, even if parent-daughter ratios have been disturbed (see below). (iii) Mantle melting produces marked fractionations of Re from Os that rapidly evolve into distinct time-integrated isotopic signatures for both melts and reservoirs. Because of these features, the Re-Os isotope system has proven particularly successful in the study of CLM. One of the objectives of this contribution is to evaluate more thoroughly the robustness of some aspects of the system.
2.1. CLM isotopic heterogeneity Once the overprint of host-magma and low-T alteration on mantle derived rocks has been properly
accounted for (see detailed leaching studies on minerals by Richardson et al. (1985), Jagoutz (1988), and Pearson et al. (1993), and review by Pearson (in press), one of the most striking features of the isotope database for CLM is its heterogeneity. Samples from both on- and off-craton settings display much greater Nd-Sr isotope variation than the oceanic mantle (Fig. 4). This heterogeneity must result from the long-term isolation of CLM from the homogenising effects of convective stirring and diffusion at the higher temperatures prevailing in convecting mantle. Heterogeneity is also evident for Pb isotopes (Menzies, 1990)) and Os isotopes (Fig. 5). Two other features are evident from these compilations of data. Firstly, although the commonly proposed origin for most CLM involves initial melt depletion (i.e., they are melting residues), in Sr-Nd isotope space, a great many samples plot in the enriched quadrant of the isotope plot. This contrasts with Os where samples predominantly have negative YOs, characteristic of an ancient, residual character. Secondly, to uniquely define CLM in Nd-Sr isotope space (or in isotope space involving Pb isotopes) is very difficult, whereas the Os isotopic composition of CLM appears to be clearly distinct from oceanic mantle, with continental crust being grossly displaced to more radiogenic values than those plotted on Fig. 5. Hence, Os appears to be a powerful tool for identifying ancient CLM and this property is being utilised in tectonically complex areas where the likely nature of the lithospheric root is unclear (e.g., Handler et al., 1997; Hassler and Shimizu, 1998). Recent analytical advances will now permit Hf isotopic characterisation of the CLM and use of potentially useful combined Hf-Os isotopes (McDonough, 1994). An idea of the length scale of isotopic variation in the CLM is probably best obtained from orogenic peridotites. Considering only the peridotites, isotopic variations that are large fractions of the total variability of oceanic mantle are present over km scale distances (Reisberg and Zindler, 1986; Zindler and Hart, 1986). When the pyroxenites are considered, isotopic variations in Nd and Pb occur over 10s of cm in most massifs that are considerably greater than those shown by oceanic mantle, (Reisberg and Zindler, 1986; Reisberg et al., 1989; Downes et al., 1991; Mukasa et al., 1991; Reisberg et al., 1991; Pearson et al., 1993; Kumar et al., 1996). Although
D.G. Pearson / Lithos 48 (1999) 171-194
175
t-
Fig. 4. Sr-Nd isotope compositions of diopsides or garnets (where garnet is the only Ca-Al-rich phase) for lithospheric peridotites compared to oceanic mantle (inset) and continental crust. Cratonic mantle peridotites solid symbols, non-cratonic peridotites open symbols. Peridotite data from references listed in Menzies (1990) plus Carlson and Irving (1994), Gunther and Jagoutz (1994, 1997), Pearson et al. (1995a,c).
not quantifiable in terms of length scale, an idea of the local CLM isotopic variability sampled by individual kimberlites bearing xenoliths, is illustrated by the 42 end unit variation in initial Nd isotope composition shown by individual garnet crystals from the Udachnaya kimberlite pipe (Jacob et al., 1998). The garnets analysed by Jacob et al. (1998) and similar garnets analysed by Pearson et al. (1995c) vary from + ve initial ~ONd values to strongly negative (to ~~ --55) and indicate the potentially complex behaviour of incompatible element-based isotope systems in CLM samples. On a single rock scale, Jagoutz (1988) found 16 ~~ units variation within different magnetic fractions of garnet from the same Tanzanian eclogite xenolith. This type of variation indicates that if melt
fractions are removed in different batches from even a single CLM sample, they have the potential to create magmas with widely varying isotopic compositions from the same source rock. Probably the largest magnitude of isotopic variability in the CLM comes from the pyroxenites and eclogites that are interspersed throughout it and sampled as xenoliths by kimberlites and other deeply derived magmas. This variability is by far the most pronounced for Os isotopes (Fig. 6), the total variation within single suites of eclogites and pyroxenites being much greater than oceanic mantle, indicating the potentially dramatic effect of incorporating such components into magmas during mantle melting. Although different isotopic systems may reflect the effects of enrichment, or depletion, the extreme
176
D.G. Pearson/Lithos 48 (1999) 171-194
nature of the isotopic variations observed is a clear indication of the general antiquity of the CLM reservoir. The quantification of how old these isotopic signatures might be is outlined below. 10
[i:...:..:t
E
Kaapvaal peridotites
6
~ "i "" " ""
4
Z
4
-
Siberian peridotites
D Low-T r:(1 H ig h -T
0
d z
2
0
f Wyoming peridotites
3 E ~o d
2
Z
1
0 6
Oceanic mantle peridotites E
4
~
E
0
Z
~
2
[! D []
3. Age constraints 3.1. Mineral and whole rock isochrons
Consideration of published diffusion coefficients e.g., Sneeringer et al. (1984), together with likely length scales of diffusion for two minerals such as diopside and garnet in a coarse grained garnet peridotite show that at lithospheric temperatures, above 1000~ diffusion will out-pace radiogenic ingrowth of daughter isotopes in most circumstances. Hence, the application of the isochron technique to examine lithospheric age will be somewhat limited (Zindler and Jagoutz, 1988). Given the effort involved in analysis, this method has yielded relatively little clear information on the chronology of lithospheric processes. For example, analysis of orthopyroxene, clinopyroxene and garnet from a low-temperature peridotite from the Premier kimberlite (Fig. 7) yielded an isochron age that is within error of the Proterozoic emplacement age of this kimberlite (Pearson et al., 1995a). Other studies that have obtained mineral isochrons defining the approximate age of kimberlite eruption include garnet lherzolites (Richardson et al., 1985; Walker et al., 1989a; Pearson et al., 1995c) and eclogites (Snyder et al., 1993; Pearson et al., 1995c). Minerals in xenoliths erupted by much younger, alkali-basalt magmas are often in, or close to, Nd isotopic equilibrium (Jagoutz et al., 1980; McDonough and McCulloch, 1987). More obvious success has been achieved where mineral isochrons have successfully been used to estimate the age of emplacement of orogenic peridotite massifs. This approach relies on garnet and pyroxene cooling past the Sm-Nd blocking temperature on emplacement of the peridotite bodies into the crust. The large difference in S m / N d fractionation between garnet and pyroxene means that even relatively young emplacement ages can be fairly precisely constrained by mineral isochrons from garnet
Ophiolites Abyssal OIB xenoliths
L
-10
Fig. 5. Histograms of Os isotopic compositions of cratonic peridotite xenoliths expressed as initial Yos values (at the time of kimberlite eruption; see Fig. 3 for calculation) compared to oceanic mantle. Data sources: Walker et al. (1989a), Luck and Allbgre (1991), Hauri et al. (1993), Carlson and Irving (1994), Snow and Reisberg (1995), Pearson et al. (1995a,c), and Pearson, unpublished.
D.G. Pearson/Lithos 48 (1999) 171-194
.~_,.~_,.~_,.~_o~'.~_,~;~_;~_;~_;~_~'._ 9
177
;_ "._ "._ ;_ ;_ ;~_'._ ;_ ;_ "._ "._ ;~_'._ ;_ "._ ;_ ;_ ;_ "._ IN_
Fig. 6. Os isotope variation (as Yo~) in various lithologies from continental roots compared to crustal rocks (data sources as in Fig. 5 plus Walker et al. (1989b), Reisberg et al. (1991), Esser and Turekian (1993), Reisberg and Lorand (1995), Pearson et al. (1995d), Burnham et al. (1998), Menzies et al. (1998), Shirey et al. (1998)).
lherzolites and garnet pyroxenites and this method has been successfully employed in this way for both Ronda and Beni Bousera peridotite massifs (Zindler et al., 1983; Reisberg et al., 1989; Pearson et al., 1995b). More ancient mineral isochrons, significantly in excess of eruption ages, have been determined for diopside-garnet pairs in cratonic peridotites (Walker et al., 1989a; Pearson et al., 1995a,c; Gunther and
]44 Fig. 7. S m - N d mineral isochron for a low-temperature garnet lherzolite from the Premier kimberlite, S. Africa. Data from Pearson et al. (1995a). Error bars are smaller than symbols, error on age is 2 sigma.
Jagoutz, 1997) and eclogites (Jagoutz, 1988; Jagoutz et al., 1984; Jacob et al., 1994). Garnet-diopside mineral isochrons from a single kimberlite pipe can vary widely in slope, giving apparent ages that are hundreds of Ma different (Fig. 8). The geochronological information provided by these isochrons is unclear. Wide variations in isochron ages from similar xenoliths in a single kimberlite are unlikely to simply represent closure ages due to post-formation cooling (Pearson et al., 1995c). In addition, even if mineral dis-equilibria was a simple function of cooling, the complex relationship between diffusion and in situ radiogenic growth (Zindler and Jagoutz, 1988) means that a given isochron "age" cannot readily be linked to any specific event. Contamination by the host kimberlite can alter peridotite mineral isochrons, complicating the situation further (Gunther and Jagoutz, 1997), but this can usually be readily identified. Gunther and Jagoutz (1997) propose that the oldest mineral isochrons in Siberian peridotites, of ca. 2 Ga, represent closure ages. Younger ages represent partial closure/re-equilibration during either lithospheric residence or during eruption. In situ trace element measurements of minerals from some of these rocks provide a different perspective. The frequent presence of both fine-scale (100 IxM) zonation, and non-equilibrium partitioning
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peridotite
~
o.52o
0.5
Fig. 8. Garnet-diopside Sm-Nd mineral isochrons for peridotites from the Mir kimberlite pipe, Siberia (McCulloch, 1989; Zhuravlev et al., 1991" Pearson et al., 1995c; Gunther and Jagoutz, 1997).
behaviour for REE between many garnets and clinopyroxenes from Siberian peridotites (Shimizu et al., 1997a) suggests disruption of isotope systematics in many peridotite minerals due to recent mineral growth. In addition, REE inter-mineral element partitioning indicates non-equilibrium between minerals in numerous instances (Shimizu et al., 1997a). These complexities make interpretation of the slope of mineral isochrons for lithospheric peridotites a subjective matter, and detailed studies of mineral zonation are required for more secure interpretation. Despite this, some xenoliths appear to give plausible mineral isochron ages such as the 1.75 _+ 0.01 Ga Sm-Nd isochron age for multiple garnet and clinopyroxene fractions from an eclogite xenolith (Jagoutz, 1988), interpreted to reflect the time of major mineral equilibration during emplacement into the lithosphere. This isochron age can be related to a regional thermal event in the Tanzanian crust. The high initial end value of 25 for the isochron implies that the source/protolith for the eclogite could be much older. It is possible that the Tanzanian eclogite studied by Jagoutz (1988) could be of crustal origin (Rudnick et al., 1998, personal communication) and that the 1.75 Ga isochron age could therefore repre-
sent a thermal effect in the crust that would have presumably manifested itself in the mantle in some way. Perhaps a more successful approach to S m / N d isochrons is to increase the diffusion length-scale by analysing minerals, e.g., diopsides, from different xenoliths that may have been separated by considerable distances in the mantle. An example of this approach is given by Deng and McDougall (1992) who analysed diopsides from a suite of spinel peridotites from Inner Mongolia. These authors obtained a good correlation on a Sm-Nd isochron diagram that yields an age of 1.64 _+ 0.01 Ga (Fig. 9). This age is taken to represent the time of their differentiation from the convecting mantle. Regional enrichment events in the mantle lithosphere can also be constrained using this approach (e.g., Carlson and Irving, 1994). Isochronous relationships have been obtained on some eclogite suites using the Re-Os system. A diamondiferous eclogite suite from Udachnaya, Siberia yields a whole rock Re-Os isochron of 2.94 _+ 0.38 Ga (Pearson et al., 1995d), whereas a diamondiferous suite from the Newlands kimberlite, S. Africa (Menzies et al., 1998) gives a whole rock
D.G. Pearson/Lithos 48 (1999) 171-194
179
spinel I
Fig. 9. Sm-Nd isochron and model ages for diopside separates from Mongolian peridotites analysed by Deng and McDougall (1992).
isochron age of 3.5 _+ 0.7 Ga. In these examples, the Os and possibly Re, reside in dispersed sulfide grains throughout the xenoliths. It appears that, at least for the Udachnaya and Newlands suite, for the length scale sampled by individual xenoliths, the Re-Os system has been effectively isolated from diffusional equilibration and has been immune enough to hostrock interaction (in contrast to Sm-Nd isotopes, e.g., Snyder et al., 1993) to preserve useful age information. The validity of the Udachnaya Re-Os isochron is supported by model age information in the case of the Udachnaya suite (see below).
3.2. U-Pb accessory phase dating Kimberlites contain a variety of accessory phases, including zircon, perovskite and baddelyite, that commonly yield U-Pb ages equivalent to the age of pipe emplacement (Davis, 1978; Heaman, 1989; Kinny et al., 1997; Scharer et al., 1997). Zircons are relatively common in highly metasomatised xeno-
liths such as MARID's and their ages can be related in some instances to enrichment events that preceded kimberlite eruption, probably earlier in the Cretaceous (Hamilton et al., 1998; Konzett et al., 1998). Some megacrystal zircons have yielded indications of very ancient derivation, in the Archaean (Kinny et al., 1989), but as they are discrete crystals it is difficult to relate this information geologically, other than to a likely lithospheric enrichment event. Better constrained information on lithospheric chronology comes from the rare crystals of zircon and monazite found within peridotite xenoliths themselves. Both these phases must be products of lithospheric enrichment and as such they relate to lithospheric modification rather than reflecting processes involved in the initial stabilisation of CLM. Nonetheless, the opportunity exists to obtain very precise estimates of enrichment ages and hence precise estimates of minimum lithosphere ages. Carlson and Irving (1994) report data from a monazite extracted from a glimmerite nodule from Montana that gives a precise
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D.G. Pearson / Lithos 48 (1999) 171-194
232Th/2~ age of 1806_+ 1 Ma. Minerals and wholerocks from the same suite scatter about a 1.8 Ga S m / N d isochron indicating that this event affected most of the lithosphere sampled, but the monazite provides a much more precise date for the enrichment. The same enrichment age has recently been confirmed by U - P b analyses of zircons (Rudnick et al., in press) from a glimmerite xenolith similar to that analysed by Carlson and Irving (1994), the average of the z i r c o n 238u/Z~ ages being 1798 + 18 Ma, but with some indication of Pb loss. In contrast to ancient mantle zircons from Montana, those extracted from a Tanzanian harzburgite (Rudnick et al., in press) are extremely young, with a m e a n 238u/Z~ age of 400 _+ 200 Ka. Thus, some of these phases must have crystallised during events closely preceding eruption (Kinny and Meyer, 1994; Rudnick et al., in press). The ancient ages recorded by monazites and zircons within both xenoliths and as megacrysts raise interesting questions about our knowledge of U - P b blocking temperatures in zircons, generally held to be ca. 800~ (Heaman and Parrish, 1991). Some of the Montana xenoliths may have very low equilibration temperatures, possibly being derived from the lower crust. However, as pointed out by Kinny et al. (1989) zircon megacrysts probably crystallise at much higher temperatures, and spend most of their existence at temperatures well above 800~ bringing into question current estimates of the blocking temperature. A > 800~ blocking temperature for U - P b diffusion in zircon is further supported by the 628 +_ age obtained for a zircon inclu12 M a 238u/Z~ sion in a diamond erupted by the Cretaceous Mbuji-Mayi kimberlite (Kinny and Meyer, 1994). If this diamond and its inclusion crystallised in the typical "diamond window" then a formation temperature above 900~ is likely. 3.3. The model age approach Because of the high ambient temperature of lithospheric mantle, and the scarcity of high U - T h accessory phases amenable to U - T h - P b dating, the most commonly used method of constraining the age of CLM is the model age approach. This technique requires a number of assumptions to be valid. One is that the rock remains a closed system since the event being dated, and another is that we have adequate
prior knowledge of the reservoir from which the rock differentiated. Because these assumptions are not always valid for the rocks being studied, model ages for mantle samples must be carefully scrutinised. For xenolith suites, the susceptibility of incompatible element-based isotope systems to disturbance by either host magma, or pre-eruption metasomatic agents, likely to have infiltrated the lithosphere over time (McKenzie, 1989), Fig. 3, means that few sample suites are likely to give consistent results. The latter process also disrupts Nd isotope systematics in orogenic peridotites (Reisberg and Zindler, 1986), with the additional complication of the likelihood of melting during ascent and emplacement into the crust that may decouple parent-daughter ratios (Pearson et al., 1993). 3.3.1. Nd isotope model ages One xenolith suite which appears to record consistent Nd-model age systematics is the Mongolian peridotite suite studied by Deng and McDougall (1992). CHUR Nd-model ages for diopsides from most of these xenoliths range from 1.1 to 1.7 Ga and cluster close to the 1.6 Ga mineral isochron (Fig. 9). Published diopside Nd isotope data from Bultfontein xenoliths also yield consistent model age results, with most samples having depleted mantle Nd model ages of ca. 1.1 Ga (Fig. 10). Although some of these samples contain garnet, isotopic evidence indicates that this phase has not equilibrated with the diopside (Gunther and Jagoutz, 1994) and contributes little to the bulk-rock Nd budget. Hence, Nd isotopes of the diopsides are a reasonable approximation of the bulk-rock Nd isotope evolution. Selection of Bulk Earth as the parent reservoir for the Bultfontein xenoliths (Fig. 10) results in more complex Nd isotope evolutionary histories, with greater variation in model ages and an overall decrease in age. The ages produced from this approach would be enrichment ages in the sense that their Nd evolution lines diverge from that of Bulk Earth at shallower angles, indicating evolution with S m / N d < Bulk Earth. This plot clearly illustrates the fundamental difficulty in deciding which parent reservoir is most appropriate for the model age calculation and the complications that this can induce. All CLM samples will be potentially subject to post-crystallisation metasomatism. Richardson et al.
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Z
Z 0.5115 0.5110 0.0
2.0
T Ga Fig. 10. Nd model ages of diopsides from Bultfontein peridotite xenoliths. Data from Richardson et al. (1985) and Gunther and Jagoutz (1994). Dashed line highlights the most aberrant sample. Although these samples contain garnet, it is usually not very radiogenic and in many cases it is unequilibrated with the diopside (Gunther and Jagoutz, 1994). The Nd budget of the rock is largely reflected by diopside in this case.
(1984) used the non-reactive, armouring properties of diamond to advantage in analysing syngenetic garnet inclusions within diamonds, with the aim of constraining the time of garnet, and hence diamond crystallisation. The garnets first analysed were of
sub-calcic composition, belonging to the harzburgitic paragenesis. Compositing of many different garnet inclusions from numerous different diamonds from the Finsch and Kimberley mines gave Nd model ages of over 3 Ga (Fig. 11). The very low eNa values of these garnet composites indicate diamond formation in ancient, enriched CLM. Analysis of two other sub-calcic garnet inclusion composites from the Premier mine (Richardson et al., 1993) produced a Bulk Earth model age older than the Earth for one sample. This result was attributed to the similar S m / N d value of the sample and reference reservoir, which combine to produce potentially large errors in estimated model ages. This potential imprecision must be borne in mind when dealing with model ages. Rigorous treatment of error propagation for model age calculations is presented by Sambridge and Lambert (1997). The ancient ages of sub-calcic garnet inclusions in diamonds (Richardson et al., 1984) is commonly quoted as some of our most reliable evidence for the antiquity of CLM, based on the premise that the diamonds have protected the garnets from subsequent interaction with mantle fluids since their encapsulation. Recently, alternative explanations have been offered for the Sm-Nd and Rb-Sr isotope systematics
r
0.0
1.0
2.0
3.0
4.0
T Ga Fig. 11. Nd model ages for sub-calcic garnet inclusions in S. African diamonds (Richardson et al., 1984, 1993).
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D.G. Pearson/Lithos 48 (1999) 171-194
of the sub-calcic garnet inclusions in diamonds analysed by Richardson et al. (1984); firstly, on the basis of the complexity of Sr isotope systematics of the garnets, that require an unusual pre-history (Pidgeon, 1989; Pearson et al., 1995c) and, secondly, on the basis of observed elemental zonation in numerous garnet inclusions in diamonds of this type (Shimizu and Sobolev, 1995; Shimizu et al., 1997b). A notable finding in this regard is that the concept that diamond is always a robust shield against postcrystallisation metasomatism is doubtful from the observation of healed fractures, observable only in cathodoluminescence, in some diamonds (Taylor et al., 1995). 3.3.2. Whole rock xenolith Os isotope model ages and system robustness Simple mixing calculations (Fig. 3) show that Os isotopic compositions are much less prone to metasomatic disturbance compared with Nd. If hostkimberlite contamination of a peridotite xenolith is considered, large mass fractions, considerably in excess of 10%, are required to obscure direct isotopic evidence of an Archaean origin (Fig. 3). Older metasomatism can be more effective in altering time-integrated Os isotope compositions due to the appreciable Re contents of most mantle melts (Fig. 3). The relatively high Re contents of mantle melts compared to peridotites means that peridotite R e / O s values are likely to reflect some component of metasomatic enrichment. This is true for both xenolith and orogenic peridotite samples. Some xenolith suites from alkali basalts show coherent Re-Os systematics that largely reflect melt depletion (Handler et al., 1997) but cratonic xenolith suites appear to be thoroughly disrupted. It is not presently clear how much of the overall disruption of R e / O s is related to mantle metasomatism and how much is due to sulfide breakdown and alteration (Lorand, 1990), but the various cratonic peridotite xenoliths that have bulk Re contents and R e / O s higher than Bulk Earth (Walker et al., 1989a; Pearson et al., 1995a) have either experienced Re addition (Fig. 12) or some early Re analyses were subject to large variable blank corrections. New data for peridotite xenoliths from the Kaapvaal craton (Carlson et al., in press) show higher percentages of xenoliths with very low Re contents, as low as 5 ppt. In contrast, more
obviously metasomatised peridotite xenoliths from the margin of the Tanzanian craton have a larger range in R e / O s with Re contents reaching over 300 ppt (Chesley et al., in press). Xenolith samples with low 187Os/188Os yet elevated Re in Fig. 12 could have experienced recent Re addition, either by the host kimberlite, or possibly carbonatite-like fluids. In addition to Re enrichment, addition of high-Os fluids may also perturb Os isotope systematics in lithospheric samples (Chesley and Rudnick, 1996). This observation is supported by the presence of abundant sulfides within deformed Tanzanian peridotites that show complex Re-Os systematics (Chesley et al., 1998). Furthermore, the observation of young diamond inclusion sulfides containing very high Os (5000 to 20,000 ppb) and Re (700 to > 2000 ppb) (Pearson et al., 1998a) illustrates the potential of sulfides crystallising from metasomatic fluids for disrupting Os isotope systematics (Fig. 12). Because of complexities due to metasomatic disturbance of various types, two model age approaches are usually adopted for the Re-Os isotope system when constraining ages in lithospheric peridotites. A "normal" model age approach (TMA ages)utilises the measured Os isotope composition and R e / O s of the sample and assumes closed system behaviour (Fig. 13). In cases where disruption of R e / O s is suspected, as in many cratonic peridotite xenoliths, the Re depletion or TRD model age is generally used. The TRD age assumes that all Re is consumed during formation of the peridotite and hence the calculation is made using a R e / O s of zero. In cases where the degree of melting is relatively low and Re is not exhausted, TMA will exceed TRD, i.e., TRD will be an underestimate of the real age (Fig. 13). In cases where the degree of melting is large, e.g., 30% or more, leading to Re exhaustion, then TMA will equal TRI~ (Fig. 13). The equations and parameters for these calculations are given in Fig. 13 and are discussed at length by Walker et al. (1989a), Pearson et al. (1995a), and Shirey and Walker (1998). The TRD age effectively gives a minimum estimate of the sample age. Any residual Re present in the sample since formation, or added by metasomatism, will decrease the calculated age of the sample. Because many kimberlite-borne xenoliths experience infiltration by their hosts, the TRD age calculation can be modified to account for post-eruption in-
D.G. Pearson / Lithos 48 (1999) 171-194
9
,,"
183
o~'" o
,"
0 o
Fig. 12. Plot of 1870S/1880S VS. ]87Re/]88Os for (A) cratonic peridotites illustrating possible field for interaction with: the host kimberlite (shaded; Pearson et al., 1995c), high Re-Os sulfides (from Pearson et al., 1998a,b), and carbonatite (Pearson et al., 1995c); (B) Pyrenean peridotites (Reisberg and Lorand, 1995). Area labelled BE is the range of values estimated for Bulk Earth by Walker et al. (1989a,b), Meisel et al. (1996), and Luck and Allbgre (1991).
growth of radiogenic Os by calculating TRD from the xenolith initial Os isotope ratio at the time of eruption (Pearson et al., 1995a). For relatively young xenoliths, unless severely metasomatised, this makes little difference to the age calculation, but for xenoliths erupted by much older magmatism, e.g., the 1.1 Ga Premier kimberlite, this correction is considerable (Fig. 13).
Some alkali basalt-borne xenolith suites, particularly those dominated by anhydrous lithologies, have relatively undisturbed R e / O s systematics, where TMA approximates to TRD within 100 Ma in some cases, e.g., Handler et al. (1997). For kimberlite-derived xenoliths, the degree of disturbance is much more variable and TMA may be considerably greater than TRD, and can often be greater than the age of the
D.G. Pearson / Lithos 48 (1999) 171-194
184
I
TRD
I
Residue 30% melting TMA = TRD ~ ,~,,~/. . . .
" ' ' .
~TRD
~~TRDerup.
1
Fig. 13. Illustration of the TMA and TRD model age approach to peridotites. TRD = 1/A * In i.Itk187Os/188OSBE(t) -- 1870S/1 88OSsample(t ))//(187Re/188OSBEt))+l] ( TMA = 1/A * In [ (187Os/188OSBE(t) -- 1870S/188 OSsample(t ) )/(187Re/188 OSBE(t) -- 187Re/188 Os sample(t)) + 1 ] where subscripts are: BE is the Bulk Earth reservoir, usually taken as chondrite; t is the time of eruption. Following Shirey and Walker (1998), parameters are: AI87Re-- 1.666.10 -11 year -l" 1870S/1880SBE at a given time is calculated from the most primitive initial 187Os/188Os for early solar system materials (IIIA irons; 0.09531) and the average lS7Re/18SOs for chondrites of 0.40186. TRD Erup shows the effect of subtracting post-emplacement radiogenic Os ingrowth of 187Os from xenoliths erupted by ancient kimberlite pipes (e.g., Premier) on the normal TRD model age. For young kimberlites the difference between TRD and TRD Erup is insignificant unless the xenolith has very high Re/Os. Inset shows hypothetical effects of various degrees of melt extraction on the Os isotope evolution of a residual peridotite.
Earth (Fig. 14). Some xenolith samples plot close to the 1" 1 TRa/TMA line on Fig. 14, in particular, the
I
..._
2.5
0.0
Fig. 14. TMA VS. TRD plot of Os model ages for kimberlite-derived xenoliths from S. Africa (Walker et al., 1989a; Pearson et al., 1995a), Siberia (Pearson et al., 1995c) and Namibia (Pearson et al., 1994). The TMA scale is truncated at 10 Ga but extends much higher.
cluster of samples close to 2 Ga, composed of samples from Namibia and the Premier kimberlite in S. Africa (see below). However, in general most cratonic samples show evidence of disturbance with TMA > TRD. In this situation, numerous samples from a given locality/region are analysed and the oldest TRO ages are taken to approximate to the age of initial lithosphere stabilisation. The spread of ages is generally interpreted to indicate later, metasomatic disturbance. One final caveat to be mindful of when interpreting Os model ages in general is that our knowledge of the present-day Bulk Earth Os isotope composition is relatively unconstrained, with estimates varying from lS7Os/lSSOs of 0.129 or higher (Martin, 1991; Meisel et al., 1996) to average carbonaceous chondrite values of ~ 0.127 (Walker and Morgan, 1989) or even lower (Luck and All~gre, 1991). In the case of cratonic xenoliths, where ages are generally quite ancient (2 Ga or more), this problem is minimised slightly due to the precisely constrained type
D.G. Pearson/Lithos 48 (1999) 171-194
IIIA iron meteorite initial Os isotope composition (Smoliar et al., 1996). However, this uncertainty remains probably the major error factor in Re-Os model ages. For example, where using a ~87Os/lgSos of 0.129 (Meisel et al., 1996) compared to average carbonaceous chondrites ( ~ 0 . 1 2 7 ) can result in model age differences of up to 300 Ma. The variability of modern day oceanic mantle has not yet been fully defined. Recently, Parkinson et al. (1998) have recorded unradiogenic 187Os/188Os as low as 0.1193 in xenoliths believed to have been derived from the oceanic mantle wedge beneath the Izu-Bonin-Mariana forearc. These peridotites give 1.2 Ga Re depletion ages using chondritic mantle Os evolution parameters. The observation of such unradiogenic Os in fragments of probably convecting oceanic mantle raises the possibility that modern oceanic mantle contains poorly mixed portions of material that has experienced older depletion events, or that the Os isotope evolution of the upper mantle has not been homogenous. Parkinson et al. (1998) suggest that subduction zones may be graveyards for ancient depleted mantle material. It remains to be seen how wide-spread this phenomena is. The implication is that if such Os isotope heterogeneity is a widespread feature of the mantle, then Re depletion model ages in peridotites must be interpreted with even more caution. However, if these environments are restricted, it might be that most lithospheric mantle generated in Archean and Proterozoic environments did not sample such heterogeneous domains, especially if plume-dominated melting is invoked for generation of ancient CLM (Herzberg, 1992; Pearson et al., 1995a). More accurately assessing the Os isotopic evolution of the Earth's mantle remains a major goal in furthering the use of Os isotopes in mantle studies. 3.3.3. Mineral Re-Os isotope model ages It is possible to try to minimise the effects of post-crystallisation R e / O s fractionation by analysing phases with very high Os and very low Re/Os, or phases capable of armouring hosts for Re and Os from later, post-magmatic interaction. Analysis of chromite from peridotite xenoliths has shown this phase to sometimes contain very high Os, but low Re contents, making it useful in estimating the initial Os isotopic composition of systems where the bulk rock
185
may have suffered substantial new Os or Re addition, reducing TRD model ages (Chesley and Rudnick, 1996; Nagler et al., 1997). It is not understood whether the high Os contents of chromites are due to micro-inclusions of sulfides, e.g., Thalhammer et al. (1990), or whether this reflects preferential partitioning of Os into the spinel structure (Capobianco and Drake, 1990). The experiments bearing on this problem have been performed in the absence of sulfide and the observation of variable Os contents in chromites point to inclusions as the dominant control. If sulfide inclusions or PGE alloys are the main cause of high Os contents within chromites, once enclosed, the chromite "armour" will be effective in protecting the sulfide grains against further interaction with metasomatic agents. Although this method may be effective in some instances, its application may be limited by the scarcity of suitable chromites. Not all chromites may contain high Os inclusions, and low-Cr spinels in xenoliths frequently contain very little ( < 1 ppb) Os (Pearson et al., 1998b). In addition, metasomatic rims are sometimes present on chromite grains which require abrading prior to dissolution (Chesley and Rudnick, 1996). For example, Chesley and Rudnick (1996) and Chesley et al. (1998) find primary, high Cr-spinels from peridotite xenoliths from Labait, Tanzania with TRD ages of 2.7-2.9 Ga., whereas secondary, Cr-rich spinels have much lower TRD ages of 2 Ga. 3.3.3.1. Regional xenolith Re-Os studies. Regions where well characterised suites of xenoliths have been analysed for Re-Os isotopic compositions, to study lithospheric evolution, include the Kaapvaal craton (Walker et al., 1989a; Pearson et al., 1995a), the Wyoming craton (Carlson and Irving, 1994), the Siberian craton (Pearson et al., 1995c), the Tanzanian craton (Chesley et al., 1998), Namibia (Pearson et al., 1994), SE Australia (Handler et al., 1997), the Eifel region (Cohen et al., 1996) and the Vitim region of southern Siberia (Pearson et al., 1998b). If histograms are plotted for TRD ages of the three cratonic peridotite data sets published so far, we see that the oldest TRD ages are between 3 and 3.5 Ga, with the Kaapvaal frequency distribution showing pronounced peaks for late- to middle-Archaean ages (Fig. 15). The oldest TRD ages for Tanzanian peridotites are from chromite separates with TRD ages as
186
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I Kaapvaal peridotites
i
]
I
1
[ Wyoming peridotites ]
Siberian peridotites
m Fig. 15. Histograms of TRD ages for peridotite xenoliths from the Kaapvaal (Walker et al., 1989a; Pearson et al., 1995a), Wyoming (Carlson and Irving, 1994), and Siberian (Pearson et al., 1995c) cratons, ages in Ga. Bushveld and Witwatersrand ages taken from Hart and Kinloch (1990). Ancient gneiss complex data are Sm-Nd ages from Carlson et al. (1983).
old as 2.9 Ga (Chesley et al., 1998, Chesley et al., in press). While bearing in mind that these ages could be minima, some striking observations can be made. For the Kaapvaal suite, the oldest TRD ages broadly match the age of the major crust building period and also are equivalent to the oldest Os ages obtained from detrital osmiridiums from the Witwatersrand basin (Hart and Kinloch, 1990) which reflect mantle extraction ages. Similar observations can be made for the Siberian and Wyoming peridotites, an Archaean age for the Siberian lithospheric keel being
supported by Archaean Re-Os model age and isochron systematics for eclogite xenoliths that are dispersed through the Siberian lithospheric keel (Pearson et al., 1995d). This coincidence of major crust building and the oldest peridotite TRD ages has been used to propose that major crust building and lithospheric stabilisation took place together from the mid- to late-Archaean (Pearson et al., 1995a). The xenolith TRD age data do not suggest any deep lithospheric mantle beneath cratons prior to ca. 3.5 Ga. In contrast, Nagler et al. (1997) found chromites within crustal ultrabasic rocks from the Zimbabwe craton that give TRD ages of up to 3.8 Ga. These ages are in excess of their emplacement ages and Nagler et al. (1997) interpreted them to reflect interaction with old lithospheric mantle during magma ascent into the crust, thus suggesting the presence of CLM beneath this region at ca. 3.8 Ga or before. The thickness of this ancient lithospheric mantle is not clear and further work is needed to investigate whether such material was present beneath the most ancient crustal fragments of other cratons at this time. The abundance and wide distribution of xenolithbearing kimberlites intruding the southern African crust, both on and off-craton, provides an opportunity to study the age relationships between cratonic and circum-cratonic mantle. Analysis of xenoliths suites intruding Proterozoic basement, off-craton in Namibia, to the West of the Kaapvaal craton, and in East Griqualand, S.E. of the craton margin (Pearson et al., 1998c) give TRD ages that range up to 2.4 Ga (Fig. 16), with no indication of Archaean ages. The oldest TRD ages for both off-craton xenolith suites match well the age of basement in each area, at ca. 2.1 Ga, showing strong coupling between mantle and crustal evolution. Even though the East Griqualand kimberlites sample mantle from within 100 km of the craton margin (Nixon et al., 1983), they are clearly distinct in age from the cratonic peridotites erupted in N. Lesotho, indicating a sharply defined craton margin in the mantle. The mid-Archaean ages of some N. Lesotho xenoliths indicate no progressive age decrease towards the edge of this part of the Kaapvaal craton and may imply a sharp tectonic contact, compatible with the abrupt age change in the lithospheric rocks. Recently, Pearson et al. (1998b) have found evidence for Proterozoic ages in peri-
D.G. Pearson/Lithos 48 (1999) 171-194
I
r
I
----1 T RD Ga
Fig. 16. Histograms comparing frequency distributions of TRD ages for peridotite xenoliths from the Kaapvaal craton (Walker et al., 1989a; Pearson et al., 1995a), Namibia (Pearson et al., 1994), and East Griqualand (Pearson et al., 1998c).
dotites from the Vitim region, surrounding the southern margin of the Siberian craton, that concur with Proterozoic crustal basement ages (Kovalenko et al., 1990). In contrast to the coherent crust-mantle relations observed above, Handler et al. (1997) found that some spinel peridotite xenoliths erupted through Paleozoic crust in SE Australia were of Proterozoic age, indicating the persistence of fragments of ancient lithosphere beneath areas of younger tectonised crust. Os isotopes provide some evidence of postArchaean lithospheric addition in the Kaapvaal craton. Samples from the Premier kimberlite, South Africa, consistently give post-Archaean ages of ca. 2 Ga (Pearson et al., 1995a; Carlson et al., in press). The Premier kimberlite erupts through the outcrop
187
ring of the Bushveld intrusion that formed close to 2 Ga and it seems likely that a major new addition to the lithosphere occurred beneath the region at this time (Richardson et al., 1993; Carlson et al., 1998). This possibility is supported by the occurrence of Bushveld erlichmenites (Hart and Kinloch, 1990) with lS7Os/188Os in the range of the Premier peridotites, that show evidence of mantle extraction at 2 Ga. The Os isotope composition of these grains presumably reflects that of the dominant component contributing to Bushveld magmatism in this period. In contrast, osmiridiums, laurites and erlichmenites from the Witwatersrand basin show evidence of grains as old as 3.3 Ga, the same age as the postulated maximum lithospheric mantle age in the main Kaapvaal craton root, south of the Limpopo belt (Pearson et al., 1995a). Further studies are in progress, aimed at characterising possible lateral variations in lithospheric age across the Kaapvaal craton to better address models of craton accretion. One significant result already evident is the presence of lithospheric peridotites with TRD ages as old as 3.6 Ga below the Limpopo Belt (Carlson et al., in press), suggesting that the Limpopo Belt is a shallow feature overlying a cratonic root that continues North, into Zimbabwe. Garnet peridotite xenoliths sampled most commonly by kimberlites have mineralogies amenable to thermobarometry, e.g., Finnerty and Boyd (1987). This potentially allows the age variation of lithospheric mantle with depth across a craton and into the surrounding circum-craton areas to be studied (Fig. 17). Plots of T~I) age vs. depth for southern African xenoliths show little correlation, with shallow, spinel-facies mantle in some pipes having similar ages to samples derived from close to the base of the lithosphere (Fig. 17). A significant result for Re-Os studies has been the finding that the hightemperature xenolith suite from kimberlites also give numerous ancient, in some cases, Archaean ages, suggesting that they have been a coherent part of the lithosphere for large periods of time (Walker et al., 1989a; Pearson et al., 1995a). This contrasts with earlier models invoked on petrological and isotopic grounds (Nixon and Boyd, 1973; Richardson et al., 1985). If the oldest xenolith TRI~ ages approximate to the actual mantle formation ages, the lack of clear age
D.G. Pearson / Lithos 48 (1999) 171-194
188
/
9
Spinel
9s
. . . . . . . . . . . . . . .
.........
2~2
.8
Graphite....... "~it"
Diamond
"
MBL
250 km
:,::7,.'S'..::;
0.9 0.5 2.4
N
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TBL Fig. 17. Schematic cross-section across southern Africa showing TRI~ ages (same data sources as Fig. 16) next to the approximate level of lithospheric derivation for selected xenoliths from individual kimberlite pipes (from mineral barometry). FL m Farm Louwrencia; J Jagersfontein; F - - Finsch; K - - Kimberley; P - - Premier; N L - - North Lesotho; R - - Ramatseliso; A - - Abbotsford.
distribution with depth in both on- and off-craton lithosphere may indicate rapid formation of lithospheric mantle keels, possibly related to plumes (Pearson et al., 1995a). While Nagler et al. (1997) suggest gradual growth of the CLM as a whole, their modelling does not preclude bursts of lithosphere generation in particular regions. In detail, the question of the time scale of generation of lithospheric keels needs to be better constrained, in particular, the proportion of CLM stabilised to different depths.
3.3.4. Re-Os systematics in orogenic peridotites Orogenic peridotites provide the opportunity to examine CLM longevity and stability in a different setting than cratons and their surroundings. As mentioned above, these massifs of peridotite are generally less isotopically disturbed than cratonic xenoliths. In addition, massif peridotites are petrologically distinct from cratonic peridotitic mantle (Boyd, 1989; Menzies and Dupuy, 1991a) and probably reflect different mantle differentiation styles. One of the most significant observations of Re-Os systematics in orogenic peridotites is the strong coherence
between major element chemistry and Os isotope systematics (Reisberg and Lorand, 1995; Reisberg et al., 1991; Burnham et al., 1998). Reisberg and Lorand (1995) observed clear correlations between 1870S/1880S and A120 3 for both Pyrenean and Ronda peridotites (Fig. 18). Re and S correlated less well. On the basis that A1 should be less easily perturbed than Re or S during mantle metasomatism, and the similar compatibility of Re to A1 during mantle melting, Reisberg and Lorand (1995) used the A120 3 -- 1 8 7 0 S / 1 8 8 0 S c o r r e l a t i o n as an isochron analogue, produced by a single melting event, to constrain the initial ~87Os/]SSOs of the peridotite suite by extrapolating A120 3 to zero. This value is then used to calculate the model age of mantle melting assuming a Re/Os of zero (essentially a TRD age). This produced a model Re depletion age of ca. 2.4 Ga for Pyrenean peridotites. Extrapolation of AlzO3-]SVOs/]S8Os correlations in peridotites assumes that both A1 and Re will be exhausted simultaneously during mantle melting. A detailed study of Pyrenean peridotites by Burnham et al. (1998) reveals good correlations for both Re
189
D.G. Pearson/Lithos 48 (1999) 171-194
1~
0.115
Fig. 18. A120 3 (wt.%) v s . 187Os/188Os correlation (R2 =0.89) for Pyrenean peridotites (Reisberg and Lorand, 1995; Burnham et al., 1998) compared to much more scattered relationship for Namibian peridotites (Pearson et al., 1994).
and S with 187Os/1880s (provided that metasomatised samples with anomalously low C u / O s are excluded), clearly showing that Re and S are both depleted before A1 in this suite, and indicating strong Re control by sulfide (Fig. 19). If estimates of initial 1870s/1880s are made by extrapolation of either Re-187Os/188Os or S - 187Os/lg8os trends, a considerably higher value is obtained than using A120 3 (Fig. 19), leading to significantly younger Re depletion ages (initial 187Os/~88Os values from Re and S correlations are within error of each other). It is unlikely that extrapolation of A120 3 --187Os/188Os correlations will be a valid procedure in any orogenic peridotites. Re appears to be controlled by sulfide in these environments and sulfide is exhausted (Lorand, 1991) well before the last significant Al-bearing phase, enstatite, with up to 5 % A12O3 in some spinel-facies peridotites. Hence, ages based on A1-Os isotope correlations will be too old. No kimberlite-borne xenolith suites have been found so far that show A120 3 correlations (e.g., see the scattered trend for Namibian peridotites in Fig. 18). This may reflect either a very different style of mantle differentiation in these samples, or, more complex disturbance of Re-Os systematics, or a combination of both. For these sample suites, frequency distributions of TRD ages from whole rocks, or model ages from chromites probably give the most meaningful results.
1870s/1880si 0.116
= 0.116
D
Fig. 19. Re-S and AI20 3 vs. 187Os/188Os plots for Pyrenean peridotite massif samples. Solid squares are "unmetasomatised" samples, open squares "metasomatised" on the basis of Cu/S values (data from Burnham et al., 1998). Note that TRD values for extrapolated initial Os isotope ratios are similar for Re and S correlations and considerably younger than the TRD model age calculated from the initial Os isotope composition obtained from the AI20 3 v s . 187Os/188Os correlation.
190
D.G. Pearson / Lithos 48 (1999) 171 - 194
Obtaining lithosphere stabilisation ages from massif peridotites appears to be possible by screening samples for metasomatism using C u / S relationships and determining initial lS7Os/lSSOs values from either Re--lS7Os/lSSOs or S-1870s/1880s correlations (Burnham et al., 1998). The Re depletion model age of 1.9 _+ 0.3 Ga obtained by Burnham et al. (1998) using this method coincides with a major period of crustal growth in the Pyrenean region at 1.86 Ga, but post-dates an older major crustal growth event at ca. 2.7 Ga, suggesting that original CLM associated with the older event has either been removed, or modified. As noted by Burnham et al. (1998), the ages given by orogenic peridotites do not provide direct information on the age of mantle deeper than the spinel-garnet transition zone in most instances. Despite the sometimes unclear relationship of orogenic peridotites to the surrounding crustal rocks, they are important groups of rocks for documenting lithospheric mantle accretion and the Os isotopic evolution of the Earth's mantle. 3.3.4.1. E c l o g i t e x e n o l i t h s . Eclogite xenoliths erupted by kimberlite pipes represent important components in the lithospheric mantle beneath many cratons and circum-cratonic regions. Their origin is currently contentious, e.g., Jacob et al. (1994), Snyder et al. (1997), but the possibility that they may represent remnants of ancient subducted oceanic lithosphere (Jagoutz et al., 1984; MacGregor and Manton, 1986; Jacob et al., 1994) means that if they can be dated accurately we may be able to constrain whether deep subduction operated in the Archaean. Dating eclogites using incompatible element-based isotope systems is as fraught with difficulty as dating peridotites (Snyder et al., 1993). Probably the most reliable "age" for eclogites using N d - S r - P b systems is the 2.76 Ga pyroxene Pb-Pb isochron reported by Jacob et al. (1994). Recently the Re-Os system has been applied to this problem (Pearson et al., 1995d; Menzies et al., 1998; Shirey et al., 1998). The relatively low Os contents of many eclogites compared to peridotites nominally makes them more susceptible to metasomatism than cratonic peridotites. However, this is counterbalanced by their radiogenic Os isotope compositions and their often very high Re/Os values that steeply diverge from the mantle evolution line (Fig. 20) such that major disturbance is required
lz9/89 a~
4.0 3.0 2.0
1.0
Fig. 20. Re-Os evolution of a suite of diamondiferous eclogites from Udachnaya, Siberia. Evolution steeply diverge from the chondritic evolution line and converge in the area between 2.5 to 3 Ga.
to dramatically affect Re-Os model age systematics. The supra-chondritic lS7Os/lSSOs of eclogites means that TMA Re-Os model ages are used rather than Tm~ ages. A suite of diamondiferous eclogite xenoliths from the Udachnaya kimberlite, Siberia, give whole-rock Re-Os model ages of 2.6 to 3.4 Ga (one sample 6.5 Ga), with Os evolution lines that intersect at ca. 2.7 Ga (Fig. 20), within error of the 2.9 _+ 0.4 Ga whole rock Re-Os isochron defined by these samples Pearson et al., 1995d). The Re-Os model ages for the Udachnaya eclogites agree well with the whole rock isochron and Pb-Pb cpx isochron for other Udachnaya diamondiferous eclogites (Jacob et al., 1994). This agreement is strong evidence that deep Archaean subduction may have operated, if these rocks are former oceanic crustal fragments, or residues from slab melting (Ireland et al., 1994). The elevated initial Os isotopic composition of the Re-Os isochron defined by Pearson et al. (1995d) may indicate a significant pre-history for the eclogite precursors.
4. Implications and summary
Some of the major discoveries made by isotopic dating of continental roots are discussed below.
D.G. Pearson/Lithos 48 (1999) 171-194
Early Sr-Nd isotope studies (e.g., Menzies and Murthy, 1980; Richardson et al., 1984) noted the record of ancient incompatible element enrichment in cratonic roots that must have evolved over billions of years to produce enriched isotopic signatures. S m - N d and Re-Os isotopic studies clearly document the fact that the CLM beneath most cratons has been isolated from the convecting mantle since at least the late- to middle-Archaean (Richardson et al., 1984; Walker et al., 1989a; Carlson and Irving, 1994; Carlson et al., in press; Pearson et al., 1995a,c; Chesley et al., in press) and possibly since the early Archaean (Nagler et al., 1997) for some cratons. No unequivocal Hadean signatures have so far been recorded. The age of CLM appears broadly similar to that of the initiation of the main crust formation events on the Siberian, Kaapvaal and Wyoming cratons. This suggests either a genetic link, or implies that large fragments of continental crust need protective deep lithospheric keels to survive. The correspondence of ancient crust underlain by ancient mantle to great depths (200 km) on craton implies long-term physical coupling of crust plus mantle in these regions that appears very robust to geological processes. Moreover, the correspondence of Proterozoic mantle beneath Proterozoic crust in several circum-cratonic regions (Pearson et al., 1994, 1998a,b) testifies to the robustness of circum-cratonic mantle in some instances, perhaps protected by very thick adjacent cratonic keels. Studies of other, predominantly spinel-peridotite xenolith suites (Cohen et al., 1996; Handler et al., 1997) and orogenic peridotite massifs (Reisberg and Lorand, 1995; Burnham et al., in press), show that continental mantle can survive for billion year periods, at least at shallow levels, in more complex tectonic settings. No convincing age structure with lithospheric depth has so far been found that gives a clear view of how cratonic roots have been constructed, perhaps indicating rapid formation of large thick keels, either by rapid arc accretion, or plume-driven differentiation. Pearson et al. (1995a) and Carlson et al. (in press) have pointed out the possibility of new lithospheric material being added to the centre of the Kaapvaal root at ca. 2 Ga during the formation of the Bushvled complex. More detailed studies are in progress, also aimed at trying to define lateral age
191
variation within cratons (Carlson et al., in press). Later additions of younger lithospheric material to the root of the Tanzanian lithosphere have been suggested by Chesley et al. (in press) on the basis of much younger (0.3 to 1.0 Ga) TRD ages obtained from xenoliths with > 120 km equilibration depths. The Re-Os system is a powerful tool for distinguishing ancient lithosphere in complex tectonic areas (Handler et al., 1997; Hassler and Shimizu, 1998) and in constraining continental keel evolution. More studies of well constrained xenolith suites and orogenic peridotites are required from different cratons and orogenic belts to further our understanding of how continents are built and evolve.
Acknowledgements The author thanks R.W. Carlson, S.B. Shirey, F.R. Boyd and P.H. Nixon for guidance on many issues relating to the material covered in this manuscript and to G. Nowell, R. Rudnick and R. Carlson for helpful reviews.
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LITHOS ELSEVIER
Lithos 48 (1999) 195-216
Nature of the mantle roots beneath the North American craton" mantle xenolith evidence from Somerset Island kimberlites S.S. Schmidberger *, D. Francis Earth and Planetary Sciences, McGilI University, 3450 University Street, Montreal, Quebec, Canada H3A 2A7
Received 23 March 1998; received in revised form 11 January 1999; accepted 12 January 1999
Abstract The recently discovered Nikos kimberlite on Somerset Island, in the Canadian Arctic, hosts an unusually well preserved suite of mantle xenoliths dominated by garnet-peridotite (lherzolite, harzburgite, dunite) showing coarse and porphyroclastic textures, with minor garnet-pyroxenite. The whole rock and mineral data for 54 Nikos xenoliths indicate a highly refractory underlying mantle with high olivine forsterite contents (ave. Fo = 92.3) and moderate to high olivine abundances (ave. 80 wt.%). These characteristics are similar to those reported for peridotites from the Archean Kaapvaal and Siberian cratons (ave. olivine Fo = 92.5), but are clearly distinct from the trend defined by oceanic peridotites and mantle xenoliths in alkaline basalts and kimberlites from post-Archean continental terranes (ave. olivine Fo = 91.0). The Nikos xenoliths yield pressures and temperatures of last equilibration between 20 and 55 kb and 650 and 1300~ and a number of the peridotite nodules appear to have equilibrated in the diamond stability field. The pressure and temperature data define a conductive paleogeotherm corresponding to a surface heat flow of 44 m W / m 2. Paleogeotherms based on xenolith data from the central Slave province of the Canadian craton require a lower surface heat flow ( ~ 40 m W / m 2) indicating a cooler geothermal regime than that beneath the Canadian Arctic. A large number of kimberlite-hosted peridotites from the Kaapvaal craton in South Africa and parts of the Siberian craton are characterized by high orthopyroxene contents (ave. Kaapvaal 32 wt.%, Siberia 20 wt.%). The calculated modal mineral assemblages for the Nikos peridotites show moderate to low contents of orthopyroxene (ave. 12 wt.%), indicating that the orthopyroxene-rich mineralogy characteristic of the Kaapvaal and Siberian cratons is not a feature of the cratonic upper mantle beneath Somerset Island. 9 1999 Elsevier Science B.V. All rights reserved. Keywords: Canadian craton; Kimberlite; Mantle root; Mantle xenolith; Peridotite; Somerset Island
1. Introduction Global seismic t o m o g r a p h y has shown that cold mantle roots with high seismic velocities underlie
* Corresponding author. E-mail: [email protected]
most Archean cratons to depths of at least 350 to 400 k m (Jordan, 1988; Grand, 1994). Kimberlite-hosted mantle xenoliths are the only direct samples of these mantle roots beneath Archean cratonic areas and thus provide constraints on their composition and structure. Studies of such xenoliths in kimberlites from South Africa have shown the highly refractory nature of the mantle roots beneath the Archean Kaapvaal
0024-4937/99/$ - see front matter 9 1999 Elsevier Science B.V. All rights reserved. PII: S0024-4937(99)00029-8
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craton, with high olivine forsterite contents ( M g / (Mg + Fe) greater than 0.91; e.g., Cox et al., 1973; Nixon et al., 1981; Boyd, 1987). A large number of these mantle nodules are characterized by high modal orthopyroxene (up to 45 wt.%; Boyd and Canil, 1997) and low M g / S i ratios (Boyd et al., 1997). The occurrence of orthopyroxene-rich mineral assemblages has also been reported for the Siberian craton (ave. O p x = 20 wt.%; Boyd et al., 1997). These compositional characteristics indicate that the deepseated mantle roots of Archean cratons are chemically distinct from the upper mantle beneath oceanic
and post-Archean crustal terranes (Boyd, 1989; Boyd et al., 1997; Shi et al., 1998). Whether the orthopyroxene-rich character reported for the mantle roots beneath the Kaapvaal and Siberian cratons is characteristic of cratonic mantle roots in general, however, has yet to be established. In this paper, we present geochemical data for a suite of mantle xenoliths from the Nikos kimberlite on Somerset Island in the Canadian Arctic. Mineral chemistry and modal mineral analysis provide constraints on the nature of the mantle roots beneath the Canadian craton and enable a comparison with the
I
/ /
,f--i
/
/
/
Fig. 1. Geological map of Somerset Island (after Steward, 1987) showing kimberlite locations. Legend: light gray, Paleozoic cover rocks; striped, Late Proterozoic cover rocks; dark gray, Precambrian basement; bold lines, normal faults.
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S.S. Schmidberger, D. Francis/Lithos 48 (1999) 195-216
compositional characteristics reported Archean Kaapvaal and Siberian cratons.
for
the
2. Geology Somerset Island is part of the Innuitian tectonic province on the northern margin of the North American continent (Trettin et al., 1972). The local stratigraphy comprises middle Proterozoic to early Paleozoic sedimentary sequences that overlie the crystalline Precambrian basement of the Boothia Uplift
(Steward, 1987). The Precambrian basement represents a northward salient of the Canadian shield that was deformed and faulted in the late Proterozoic to early Paleozoic, possibly related to compression associated with the Caledonian Orogeny in northern Europe and Greenland (Okulitch and Trettin, 1991). Granulite facies gneisses from the crystalline basement of the Boothia Uplift yield zircon dates between 2.5 and 2.2 Ga for their protoliths, while S m - N d model ages of 3.0 to 2.2 Ga suggest the presence of reworked Archean crustal material in these gneisses (Frisch and Hunt, 1993).
Table 1 Major and trace element analyses of kimberlites The major and trace elements were analyzed by X-ray fluorescence analysis at McGill University with a Philips PW 2400 spectrometer. The major elements and Ba, V, Co, Cr, and Ni were analyzed on fused discs using a c~-coefficient technique with an analytical precision of less than 0.06 wt.% (1 SD). Total Fe is given as FeO. Other trace elements were analyzed on pressed pellets using a Rh Kb Compton scatter matrix correction with an analytical precision of 5%. Sample Rock
NK3-K1 Kimberlite
Major elements in wt. % SiO 2 19.85 TiO 2 2.06 A120 3 2.20 FeO 7.06 MnO 0.14 MgO 19.97 CaO 21.58 Na20 0.19 K20 1.00 P205 0.64 H20 1.43 CO 2 22.78 Total 98.90 CaCO3 38.52 Trace elements in ppm Ba 2694 Rb 72 Sr 1299 Th 18.2 U 0 Zr 139 Nb 166 Y 9.3 Cr 1366 Ni 672 Co 52 V 153 Sc 28
NK3-K2 Kimberlite
NK3-K3 Kimberlite
NK3-K4 Kimberlite
NK3-K5 Kimberlite
25.05 2.15 2.81 8.13 0.16 26.12 15.34 0.14 0.66 0.64 5.00 12.68 98.88 27.38
23.21 1.62 2.03 6.99 0.15 23.07 17.57 0.06 0.55 0.92 5.07 17.10 98.35 31.36
22.72 1.73 1.90 7.32 0.14 23.03 18.90 0.09 0.30 0.83 8.55 13.19 98.71 30.00
23.62 1.51 1.64 7.26 0.15 25.59 16.61 0.10 0.59 0.85 0.00 20.96 98.88 29.65
2231 46 1072 17.4 1.4 162 169 11.3 1586 817 64 151 16
2972 30 1647 22.3 0 163 198 11.8 1721 861 59 136 15
3007 22 2233 23.1 0 146 190 9.5 1526 855 64 140 19
2445 38 1418 18.7 0 147 179 10.4 1362 917 65 126 23
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s.s. Schmidberger, D. Francis / Lithos 48 (1999) 195-216
Other kimberlite occurrences on Somerset Island and their mantle xenoliths have been previously de-
scribed (e.g., Mitchell, 1977; Jago and Mitchell, 1987; Fig. 1). The kimberlite pipes were emplaced
(a)
Fig. 2. Photomicrographs of xenoliths from the Nikos kimberlite showing (a) coarse texture in peridotite, (b) porphyroclastic texture in peridotite, and (c) coarse texture in pyroxenite. Upper photographs with plane polars, lower photographs with crossed polars.
S.S. Schmidbergrr, D. Francis/Lithos 48 (19991 195-216
199
Fig. 2 (continued).
during the Cretaceous along a major fault system related to the development of the Boothia Uplift (Mitchell, 1975). U-Pb dates on perovskite indicate
kimberlite formation ages between 88 and 105 Ma (Heaman, 1989; Smith et al., 1989). Kimberlite emplacement coincides with major compressive defor-
200
S.S. Schmidberger, D. Francis /Lithos 48 (1999) 195-216
Fig. 2 (continued).
mation in the southeastern Arctic Islands induced by rifting that preceded sea-floor spreading in the Labrador Sea and Baffin Bay and the associated
rotation of Greenland away from the North American continent that began in the early Cretaceous (Okulitch and Trettin, 1991). The timing of kimber-
S.S. Schmidberger, D. Francis / Lithos 48 (1999) 195-216
lite emplacement on Somerset Island coincides with a world wide period of kimberlite activity in the Cretaceous (Nixon, 1987).
3. Kimberlite host
Previous field work and petrographic studies have shown that the majority of the Somerset kimberlites are brecciated diatremes or hypabyssal roots zones (Mitchell and Meyer, 1980), consisting of fragments of kimberlite host, country rock, and mantle-derived xenocrysts. Establishing the compositions of the liquids for such kimberlites is problematic because of the difficulty of accounting for the country rocks and xenocrysts present in the kimberlite breccia. Magmatic kimberlite, containing less than 15 vol.% of clasts larger than 4 mm, is rare on Somerset Island (Mitchell and Meyer, 1980; Mitchell, 1986). The recently discovered Nikos kimberlite (Pell, 1993) outcrops along a steep cliff on the eastern coast of Somerset Island (73~ and 90~ Fig. 1) and consists of three individual pipes that were sampled during the summer of 1996. Petrographic observations of the kimberlite samples indicate that one of these pipes (NK3) is characterized by a non-brecciated, magmatic texture. The magmatic kimberlite exhibits a porphyritic microcrystalline texture containing phenocrysts of olivine, phlogopite, and spinel in a very fine-grained carbonaterich matrix. Two generations of olivine can be identified, rounded olivine phenocrysts ( > 1 mm) that most probably crystallized prior to the intrusion of the kimberlite and small euhedral to subhedral olivines ( < 1 mm). Garnet megacrysts (up to 5 mm) may represent xenocrysts derived by fragmentation of garnet peridotite xenoliths or be a liquidus phase at high pressure in the kimberlite magma (Mitchell, 1978; Mitchell and Meyer, 1980). The microcrystalline groundmass consists of calcite, serpentine, perovskite, and apatite. Calcite occurs as aggregates of tabular euhedral crystals or as sub-parallel laths, which are deflected around large olivine crystals, indicating their primary magmatic nature. Serpentine occurs as spherical microcrystalline segregations and as an alteration product of olivine marginally replacing these phenocrysts.
201
The whole rock analyses of kimberlite NK3 (n = 5) indicate an extremely carbonate rich composition, with CaCO 3 contents between 27 and 39 wt.% (ave. 31 wt.%), low SiO 2 abundances (20-25 wt.%), and a strong enrichment in incompatible trace element contents ( B a - 2 2 0 0 - 3 0 0 0 ppm; S r = 1100-2200 ppm; Table 1). The high carbonate contents of the kimberlitic liquid indicate compositions intermediate between kimberlite (ave. CaCO3 = 14 wt.%; Mitchell, 1986; Ringwood et al., 1992) and carbonatite (CaCO 3 = 50-80 wt.%; Woolley and Kempe, 1989), in keeping with recent experimental studies that have shown that a continuous spectrum of melt compositions from carbonatite to kimberlite can be produced by increasing degree of melting at pressures on the order of 60 kb (Dalton and Presnall, 1997).
4. Mantle xenoliths
The Nikos kimberlite hosts an unusually well preserved suite of large (10-20 cm) mantle-derived xenoliths, 54 of which were collected and analyzed in this study. The Nikos xenolith suite is dominated by garnet-peridotites (about 90%) comprising lherzolites, harzburgites, and lesser dunites. Also present are small numbers of pyroxene-rich nodules that are classified as garnet-pyroxenite rather than eclogite (Dawson, 1980; Hatton and Gurney, 1987) because of their low modal garnet (10-15 wt. %) and A12 O3 contents (5-10 wt.%), and the low jadeite component in clinopyroxene (Na20 < 1.6 wt.%). The majority of the peridotite nodules have textures ranging from coarse equant to coarse tabular (Fig. 2a), based on the textural classification for mantle-derived rocks of Harte (1977). Large subhedral crystals of well preserved and relatively strainfree olivine (up to 10 mm) and orthopyroxene (ave. 2-5 mm) dominate the mineral assemblage. Clinopyroxene occurs as emerald green anhedral crystals with smooth grain boundaries (1-3 mm). Purple red garnet occurs as large porphyroclasts (ave. 5 mm) and is variably surrounded by kelyphitic rims of phlogopite and small spinel crystals. A small number of peridotite xenoliths exhibit porphyroclastic textures ( < 10%; Fig. 2b), characterized by large porphyroclasts (up to 10 mm) of strained olivine
Table 2 Representative analyses of bulk xenoliths LOI: loss on ignition. Total Fe is given as FeO and Mg # = M g / ( M g +Fe). Lherz: lherzolite, Harz: harzburgite, Pyrox: pyroxenite, porphyr: porphyroclastic. Modes were calculated using the whole rock compositions and a high-pressure peridotite norm recipe. Analytical procedure as in Table 1. Sample Rock Texture
NK1-5 Lherz coarse
NK2-3 Lherz coarse
Major elements in wt. %. SiO 2 42.28 42.46 TiO 2 0.07 0.09 A120 3 1.95 2.90 FeO 7.74 7.21 MnO 0.12 0.13 MgO 41.49 41.87 CaO 1.44 1.98 Na20 0.06 0.15 K20 0.08 0.19 P2 05 0.01 0.02 Cr20 3 0.85 1.04 NiO 0.30 0.31 LOI 3.50 1.77 Total 99.89 100.13 Mg # 0.905 0.912
NK2-10 Lherz coarse
NK3-20 Lherz coarse
NK3-25 Lherz coarse
NK1-2 Harz coarse
NK1-3 Harz coarse
NK1-6 Harz coarse
NK1-7 Harz porphyr
NK1-17 Harz coarse
NK1-18 Harz coarse
NK2-7 Pyrox coarse
NK3-1 Pyrox coarse
NK3-17 Pyrox coarse
41.54 0.03 1.25 7.43 0.11 45.53 1.22 0.07 0.05 0.04 0.53 0.32 1.84 99.97 0.916
43.19 0.10 3.21 7.58 0.13 37.50 3.21 0.35 0.14 0.02 0.67 0.28 3.30 99.68 0.898
41.42 0.21 4.28 7.41 0.13 39.46 2.13 0.09 0.52 0.04 1.09 0.25 2.81 99.85 0.905
41.77 0.03 1.31 7.39 0.11 44.44 1.11 0.10 0.03 0.02 0.52 0.32 2.71 99.85 0.915
40.81 0.03 1.33 7.42 0.11 44.03 0.82 0.01 0.04 0.02 0.60 0.33 3.89 99.42 0.914
40.37 0.02 1.25 7.10 0.10 43.41 0.41 0.00 0.03 0.01 0.68 0.33 5.85 99.57 0.916
41.24 0.03 1.57 7.06 0.10 43.55 0.72 0.00 0.02 0.01 0.70 0.31 4.31 99.62 0.917
41.22 0.01 0.81 7.05 0.11 45.40 0.49 0.08 0.03 0.03 0.68 0.31 3.88 100.10 0.920
42.94 0.03 0.79 6.97 0.10 44.93 0.54 0.05 0.03 0.01 0.51 0.35 2.75 100.00 0.920
48.95 0.13 8.05 4.93 0.18 23.69 9.45 0.90 0.11 0.01 1.69 0.38 1.90 100.37 0.895
50.97 0.18 4.71 6.04 0.15 19.44 16.46 0.46 0.11 0.02 0.79 0.09 0.52 99.94 0.852
50.15 0.18 4.89 6.01 0.15 19.48 16.91 0.28 0.09 0.01 0.73 0.09 0.58 99.55 0.852
4~
I
t~
Calculated Oliv Opx Cpx Garn
modes in wt. %. 71.0 69.0 18.1 15.8 6.1 8.6 4.8 6.6
87.8 4.2 5.1 3.0
56.9 22.5 14.2 6.4
60.9 19.5 9.8 9.9
83.5 8.8 4.7 3.0
84.1 9.1 3.4 3.5
82.5 12.3 1.7 3.5
79.2 13.8 2.9 4.1
87.9 7.7 2.3 2.2
80.6 15.1 2.3 2.0
4.4 41.0 39.2 15.4
0 23.6 67.0 9.4
0 22.2 67.6 10.3
S.S. Schmidberger, D. Francis/Lithos 48 (1999) 195-216
with undulose extinction surrounded by strain-free olivine neoblasts (0.1-0.5 mm) with tabular to polygonal equant habits. In these rocks, orthopyroxene occurs as porphyroclasts (2-5 mm) with irregular habit, and clinopyroxene and garnet are optically similar to those found in the coarse peridotites. The pyroxenite nodules show coarse equant textures (Fig. 2c) with large subhedral crystals (2-5 mm) of clinopyroxene and orthopyroxene, and smaller crystals (1 mm) of pale red garnet with a polygonal habit.
5. Analytical data 5.1. Whole rock chemistry Whole rock analyses indicate that the Nikos peridotites are strongly depleted in fusible elements such as Fe, Ti, Ca, A1, and Na compared to estimated primitive mantle compositions (McDonough, 1990" Table 2). The whole rock magnesium numbers (Mg # = M g / ( M g + Fe); Fig. 3) for the lherzolites (ave.
203
Mg # = 0.91), harzburgites (ave. Mg # = 0.92), and dunites (ave. Mg # = 0.92) indicate a shift toward higher values for the harzburgites and dunites consistent with their larger degree of depletion. The Nikos nodules are significantly more refractory than oceanic peridotites and mantle xenoliths derived from beneath post-Archean continental crust (Mg # = 0.890.91; Boyd, 1989). There is a good negative correlation between A1 and Ca, and Mg (Fig. 4a,b), and a rough positive correlation between Ni and Mg (Fig. 4c). Compared to spinel-lherzolites from the Canadian Cordillera (Shi et al., 1998), the late Phanerozoic orogenic belt that borders the western margin of the Canadian shield, the Nikos peridotites contain much lower A120 3 contents (Fig. 5) and higher Mg#s (Fig. 8). This indicates that the mantle roots beneath the Canadian shield are significantly more refractory than the lithospheric mantle beneath the Phanerozoic Cordillera. Spinel-harzburgites occurring in some of the Cordilleran xenolith suites from the southwestern Yukon and northern British Columbia, however, exhibit A1203 contents indistinguishable from those for the Nikos xenoliths (Fig. 5).
Fig. 3. Histogram of whole rock magnesium numbers (Mg # = Mg/(Mg + Fe)) of peridotites, and high-Mg and low-Mg pyroxenites.
204
S.S. Schmidberger, D. Francis / Lithos 48 (1999) 195-216
lherzolites and 0.92 to 0.93 (ave. 0.925) in the harzburgites, indicating higher values for the more refractory nodules (Table 3 for representative analyses). Individual olivine crystals do not exhibit chemical zoning and are homogeneous on thin section scale. There is no compositional difference between the olivines of the coarse and porphyroclastic peridotites. The Somerset olivine Mg#s (ave. 0.923) are similar to those for peridotites from the Kaapvaal and Siberian cratons (ave. 0.925), but somewhat higher than the values reported for the Jericho (ave. 0.913; Kopylova et al., 1999) and the Torrie xeno-
|
|
|
9
b
@
9 9 9
0.3 60 |
i Bi-modal
Fig. 4. Variations of major elements (a) A1203, (b) CaO, (c) NiO with MgO (wt.%) of Nikos peridotites. Primitive mantle composition after McDonough (1990). These refractory harzburgites are associated with a seismically detected hot mantle anomaly (Shi et al., 1998). The pyroxenites, on the other hand, show a large variation in Mg#s. The pyroxenites having high Mg#s between 0.88 and 0.92 (high-Mg pyroxenites) overlap those of the peridotites. The Mg#s for the other group of pyroxenites, however, are significantly lower (Mg # = 0.84-0.85; low-Mg pyroxenites), indicating that they are more Fe-rich than the coexisting peridotite xenoliths (Fig. 3). 5.2. Mineral chemistry 5.2.1. Olivine Olivine in the peridotites is magnesium-rich and ranges in Mg # from 0.91 to 0.93 (ave. 0.921) in the
i i i i i i
U
,i
i
i i I
1 I I
!Cordillera ~ Uni-moda!
i 10
Fig. 5. Histogram of A1203 content (wt.%) of Nikos peridotites compared to data for the bi-modal northern Cordilleran spinelharzburgite and spinel-lherzolite suite and the uni-modal southern Cordilleran spinel-lherzolite suite (Shi et al., 1998). Primitive mantle composition after McDonough (1990).
205
S.S. Schmidberger, D. Francis / Lithos 48 (1999) 195-216
Table 3 Representative analyses of olivines The mineral compositions (wt.%) were determined using a JEOL 8900L electron microprobe at McGill University. The mineral phases were analyzed by wavelength dispersive spectroscopy with an acceleration voltage of 20 kV, beam current of 20 nA, and a beam diameter of 5 txm. On-peak counting times were 20 s, and the background counting times 10 s. Sample Rock
NK1-5 Lherz
NK2-3 Lherz
NK2-10 Lherz
NK3-20 Lherz
NK3-25 Lherz
NK1-2 Harz
NK1-3 Harz
NK1-6 Harz
NK1-7 Harz
NKl-17 Harz
NKI-18 Harz
SiO 2
40.75 8.15 0.12 49.43 0.04 0.04 0.35 98.94 0.915
41.43 7.07 0.10 51.04 0.01 0.00 0.40 100.08 0.928
40.78 7.43 0.10 50.83 0.01 0.01 0.36 99.56 0.924
41.26 8.59 0.09 50.27 0.01 0.00 0.38 100.63 0.912
41.29 8.47 0.11 50.48 0.05 0.02 0.36 100.83 0.914
41.02 7.48 0.10 50.20 0.01 0.01 0.35 99.19 0.923
41.03 7.37 0.10 50.43 0.01 0.01 0.35 99.33 0.924
40.81 7.15 0.10 51.30 0.02 0.02 0.38 99.79 0.927
41.01 7.35 0.10 50.67 0.05 0.04 0.37 99.62 0.925
40.94 7.09 0.10 51.04 0.02 0.01 0.36 99.58 0.928
41.26 7.21 0.10 50.88 0.03 0.04 0.43 99.99 0.926
0.993 0.151 0.002 1.845 0.000 0.000 0.007 3.992
1.000 0.174 0.002 1.816 0.000 0.000 0.007 4.000
0.998 0.171 0.002 1.819 0.001 0.000 0.007 3.997
1.004 0.153 0.002 1.832 0.000 0.000 0.007 4.004
1.002 0.151 0.002 1.837 0.000 0.000 0.007 4.002
0.989 0.145 0.002 1.854 0.000 0.000 0.007 3.989
1.004 0.150 0.002 1.835 0.001 0.001 0.007 3.999<
