The Crust: Treatise on Geochemistry

Executive Editors' Foreword H. D. Holland Harvard University, Cambridge, MAo USA

00' K. K. Turekian Yale University. New Haven,

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3.10.2.5 U–Th–Pb Dating by ICP-MS In the early 1980s quadrupole mass spectrometers using argon plasma as the ionization source, i.e., inductively coupled plasma mass spectrometers (ICP-MS), were developed. Although these instruments were designed primarily for measuring the concentrations of trace elements, many studies have employed them for U – Pb dating. 206Pbp/238U, 207Pbp/235U, and 208Pbp/232Th dates can be determined, in addition to the more simply determined 207Pb– 206Pb dates,

Figure 1 U – Pb Concordia diagram showing the results of an analysis of a detrital zircon crystal by ion microprobe (SHRIMP) followed by analysis of the same crystal using TIMS. Both error ellipses are plotted at 2s. The best estimate of the age of crystallization of the zircon is identical for both techniques; however, the TIMS analysis is an order of magnitude more precise than that obtained using the ion microprobe (source Samson et al., 2003).

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using laser-ablation ICP-MS techniques. The obvious appeal of U – Th – Pb dating by laser ablation ICP-MS is the elimination of the need for ultra-low blank dissolution and U – Th– Pb separation procedures, the speed of the analysis (,10 min per analysis), and the possibility of in situ analysis. However, there are several analytical obstacles to obtaining accurate U – Th– Pb dates via laser ablation that must be overcome. One of the most difficult problems is that there is significant elemental (i.e., U/Pb and Th/Pb) fractionation that occurs during laser ablation. That is, measured Pb/U ratios are lower than actual ratios by tens of percent, and this effect is variable with ablation time (see figure 4 in Horn et al., 2000). A second effect, common to all mass spectrometric measurements, is that there is an instrumental mass bias, or discrimination. This bias is several times higher than the bias that occurs during TIMS measurements and thus would be a significant source of error if not corrected. A third potential difficulty is determining the amount of common lead in an analysis, as the argon gas used in ICP-MS contains enough mercury to cause isobaric interference of 204Pb from 204Hg. However, continued improvements are being made as this technique evolves (see below) and it may begin to approach ID-TIMS analysis in the future, at least for accessory minerals with relatively high radiogenic lead contents. Early attempts to directly date zircon crystals using ICP-MS techniques involved the use of Nd – YAG lasers, operating at a 1,064 nm wavelength, to ablate the zircon crystals (Feng et al., 1993; Fryer et al., 1993). Because of significant variations of U/Pb isotopic ratios, these early laser-ablation studies concentrated on determining 207 Pb – 206Pb dates, which yielded precision between 0.5 – 6.0% (e.g., Fryer et al., 1993; Jackson et al., 1996). However, elemental fractionation during laser ablation decreases with decreasing wavelength (Geersten et al., 1994), and thus by quadrupling the frequency of Nd –YAG lasers (266 nm wavelength), or using gas-based lasers operating in the deep UV (such as Ar – F excimer lasers which produce light at 193 nm (e.g., Eggins et al., 1998)), more reproducible U/Pb ratios could be measured compared to the earlier analyses using larger wavelengths. To counterbalance the effects of laser-induced elemental fractionation, as well as minimize temporal variations, many workers analyze externally calibrated standard minerals and unknowns under identical operating conditions, and then apply correction factors to the unknowns (e.g., Ferna´ndez-Sua´rez et al., 1999; Knudsen et al., 2001). In this respect, the technique shares strong similarities with ion microprobe U –Th – Pb age determinations. A major difference between the

326

Ages and Growth of the Continental Crust from Radiogenic Isotopes

techniques, however, is that the volume of the pit excavated by a laser is much larger than the spot produced by the ion microprobe (Figure 2), and thus laser ablation must be considered a destructive technique. Precision of 206Pb/238U dates using externalstandard correction methods with frequency quadrupled lasers is partly dependent on sufficiently high uranium and lead concentrations, but typical values appear to be in the range of several percent. Even higher reproducibility of U/Pb ratios (,2% for 206Pb/238U of a Paleozoic zircon standard) from ablated zircons was reported by Horn et al. (2000). These workers used an excimer laser (193 nm wavelength), which presumably results in less laser-induced elemental fractionation than studies employing 266 nm wavelength lasers. In addition, Horn et al. (2000) combined the laser ablation with solution nebulization of known quantities of Tl and enriched 235U. Comparing the measured 205 Tl/235U with the known ratio allowed for a correction factor for instrumental mass bias to be used on the measured 206Pb/238U ratio. By comparing the measured 203Tl/205TI ratio of the introduced Tl with its natural ratio, a correction factor was applied to the measured lead isotopic ratios to further increase accuracy, an approach first described by Longerich et al. (1987). Elemental fractionation produced at the laser-ablation

site is still significant, however, and a correction must be applied to obtain accurate data. Horn et al. (2000) demonstrated that such fractionation is a function of pit geometry, thus allowing corrections to be made by establishing an empirical correction curve using a standard zircon. In the 1990s, mass spectrometers with a magnetic sector and a full array of Faraday detectors were coupled with a plasma source. These multi collector instruments (MC – ICP – MS) produce the same flat-topped peaks produced by TIMS and thus are capable of higher-precision isotope ratio measurements than quadrupole-type instruments. There is considerable interest in using these newgeneration instruments for U – Th –Pb dating by directly analyzing zircon and monazite crystals via laser ablation, largely following the techniques originally developed for quadrupole instruments. However, because MC –ICP – MS is such a new technique, few U –Pb geochronological studies have so far been published discussing the results obtainable using these multicollector instruments (Machado and Simonetti, 2001). Based on the limited data currently available, it appears that typical 2s uncertainties for U – Pb dates on zircon using multicollector instruments (e.g., 1 – 2%) are better, but not yet substantially so, than those obtainable using quadrupole instruments. However, use of multicollector ICP – MS for U – Pb dating is in its infancy and thus the technique may hold considerable potential as a geochronological tool.

3.10.2.6 U–Th–Pb Dating of Monazite Using Only Uranium, Thorium, and Lead Concentrations

Figure 2 Size of a typical pit produced in an accessory mineral using an ion microprobe during an 18 min analytical run (from Stern, 1997) compared to the size of an ablation crater made from a single pulse of an excimer laser (from Horn et al., 2000). Bottom drawings show generalized cross-sections of the spots made from the two techniques. Note the considerably smaller volume of mineral excavated during the ion microprobe analysis.

The mineral monazite, a uranium- and thorium-rich phosphate of a rare-earth element (REE), is a common accessory mineral in a variety of felsic igneous rocks and is a common trace constituent in many metamorphic rocks, particularly metapelites. Because of its high uranium and thorium content, and fairly low common lead content, it has often been dated using ID-TIMS and IMP techniques. In the early 1990s the potential of determining the age of monazite using an electron microprobe (EMP) was investigated (Suzuki and Adachi, 1991; Montel et al., 1994; Suzuki and Adachi, 1994). The so-called “chemical age” of a monazite crystal can be determined solely by measuring its thorium, uranium, and lead contents (no isotopic measurement) if the amount of common lead is negligible and no post-crystallization loss or gain of uranium, thorium and lead occurred (see review in Montel et al., 1996). If these conditions are met, then within ,100 – 200 million years

Determination of Ages of Metamorphism p

the amount of accumulated Pb in a typical monazite is high enough that it can be measured accurately and an age calculated. Precision of the chemical age is largely governed by lead content (as is true with all in situ techniques) and thus is largely age dependent, but 2s precisions of 10 – 20 Myr are now obtainable for very lead-rich (.2,000 ppm) crystals, with errors around double those values for monazite with lower lead contents (Williams and Jercinovic, 2002). The main advantage of the EMP technique is the excellent spatial resolution that can be obtained; in situ analysis of monazites as small as 5 mm can be obtained from polished thin sections (Montel et al., 1996). This feature is of limited utility in dating relatively large (radius ¼ 30 mm) monazites with simple histories, but is invaluable for studies where in situ analysis is critical, such as determining the age of monazites that occur as inclusions in porphyroblasts (e.g., Williams et al., 1999; Montel et al., 2000). Employing such high spatial resolution must be used with considerable caution, however, as accurate ages can only be obtained if no lead diffusion has occurred within the restricted region of the crystal being analyzed. The current emphases of the technique are to constrain the timing of multiple metamorphic events, to determine directly the timing of deformational events, and to provide links between metamorphism and deformation (see review by Williams and Jercinovic, 2002). The main limitation of chemical dating of monazite by EMP, aside from the necessity of perfect closed-system behavior of the region of the mineral being analyzed, is the detection limit of lead, typically a few hundred ppm for most instruments. This limitation usually precludes the analysis of monazites younger than ,100– 200 Ma. The feasibility of determining chemical ages of young monazites containing only a few tens of ppm lead was demonstrated by Cheburkin et al. (1997) by using a newly designed X-ray fluorescence microprobe. Improvements on the original instrument design allow for chemical dating of monazite as young as 15 Ma and as small as 50 mm (Engi et al., 2002); however, the monazites could not be measured in situ. In a companion study, Scherrer et al. (2002) determined chemical ages of monazites that were first optically examined in thin section, thus still preserving full textural context of the analysis, but were then removed from the thin sections by drilling with a diamond microdrill. Although this is a labor-intensive procedure compared to EMP dating, and still cannot be done on very small monazites, an X-ray microprobe age of 55.3 ^ 2 Ma was determined for a 54 ^ 1 Ma monazite (208Pb/232Th TIMS date), demonstrating the much higher precision of the X-ray microprobe technique compared to the EMP technique.

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3.10.3 DETERMINATION OF AGES OF METAMORPHISM From the very beginning of the pioneering days of geochronology, it was noted that different minerals from a single rock had different apparent ages, suggesting that different minerals retained different proportions of radiogenic daughter nuclides (e.g., Wetherill et al., 1955), thus setting the stage for future thermochronologic studies. By 1959, the use of radioactive decay schemes to estimate the timing of metamorphism had been specifically discussed (Compston and Jeffery, 1959; Tilton et al., 1959). Subsequently, a large number of papers were published involving Rb –Sr and K – Ar dating of different minerals within metamorphic rocks, establishing the beginning of a database of the history of continental crustal deformation. With the introduction of the 40 Ar/39Ar technique (Merrihue and Turner, 1966) and the seminal discussion of the concept of mineral closure temperature (Dodson, 1973) the field of thermochronology became firmly established and its impact on our understanding of the tectonic evolution of orogenic belts has been profound. Because Chapter 3.08 is devoted to the discussion of crustal metamorphism, only a brief review of current thermochronologic techniques and their applications to continental crustal evolution is given here. Because the U –Th – Pb system is discussed in detail above, a separate section on its application specifically to constraining the timing of metamorphism is beyond the scope of this chapter. Reviews of U – Pb geochronology applied to metamorphic studies can be found in Heaman and Parrish (1991) and Mezger and Krogstad (1997). Recent discussions of the formation of metamorphic zircon domains and interpretation of geochronologic data can be found in Fraser et al. (1997) and Bingen et al. (2001), and reference therein.

3.10.3.1 40

40

Ar/39Ar Thermochronology

Ar/39Ar dating is a variation of conventional K – Ar technique in that potassium-bearing samples are irradiated with neutrons to produce 39 Ar from 39K, thereby eliminating the need for separate measurements of potassium and argon on two separate aliquots of a sample (see McDougall and Harrison, 1999 for a detailed review). In the 1970s most studies of metamorphosed crustal regions utilized the technique in a similar fashion to previous K – Ar studies, i.e., determining 40 Ar/39Ar dates of different metamorphic minerals and inferring the cooling history based on estimates of the argon closure temperature in the mineral (e.g., Lanphere and Albee, 1974; Dallmeyer, 1975; Dallmeyer et al., 1975, and

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Ages and Growth of the Continental Crust from Radiogenic Isotopes

many others). In these studies milligram-sized mineral separates were step-heated in a furnace and the argon gas released during each temperature step was isotopically analyzed. Lasers were used to heat smaller samples, first on lunar rocks (e.g., Megrue, 1973), and subsequently on terrestrial samples (e.g., York et al., 1981; Maluski and Schaeffer, 1982; Sutter and Hartung, 1984), although most of these early studies produced only total-fusion 40Ar/39Ar dates. Layer et al. (1987) demonstrated that detailed age spectra could be determined from single hornblende and biotite crystals using a defocused continuous laser beam. The advantage of laser microprobe 40Ar/39Ar analyses is the ability of in situ analysis. Spatial resolution of 50 –100 mm can be achieved using lasers in the visible and near-infrared wavelengths, although these are best for minerals that are strong absorbers of such wavelengths such as biotite, phlogopite, and hornblende (see review by Kelley, 1995). With the employment of UV lasers in 40Ar/39Ar thermochronology spatial resolution increased considerably (,10 mm width) as well the ability to analyze most silicate minerals, including white mica and feldspar that are poor absorbers of higher-wavelength energy (Kelley et al., 1994). Important applications of this technique to crustal metamorphic studies include the direct dating of deformation fabrics (e.g., Reddy et al., 1996), dating different portions of single P– T paths (e.g., DiVincenzo et al., 2001) and dating of mineral inclusions in porphyroblasts (Kelley et al., 1997). Further discussions of the modern applications of 40Ar/39Ar dating to constrain the timing of metamorphic events are given by Hodges in Chapter 3.08.

3.10.3.2 Rb–Sr Dating It was recognized early on that the Rb – Sr system was particularly useful for constraining the timing of metamorphic events because of the significant degree of rubidium and strontium diffusion that occurs between minerals during metamorphism. By constructing Rb – Sr mineral isochrons the timing of diffusion (i.e., metamorphism) can be determined, assuming complete isotopic re-equilibration occurred during a discrete metamorphic event and the system remained closed to any further disturbance (e.g., Fairbairn et al., 1961). Important early work in contact metamorphic zones constrained the behavior of Rb – Sr in mineral systems (Hart, 1964; Hanson and Gast, 1967). The Rb – Sr method continues to play a very important role in studies of the deformational and metamorphic history of crustal regions, but the focus has shifted towards determining the ages of minerals within a

well-defined textural context to better interpret the significance of the constructed isochrons. Some recent examples are given below. Because of its common occurrence in metamorphic rocks, garnet separates have been an obvious choice of one of the components to be incorporated in mineral Rb – Sr isochrons in metamorphic studies (see above). Building on that previous work, Christensen et al. (1989) sliced large (3 cm) single garnets into separate pieces for Rb –Sr isotopic analysis along with the rock matrix between the garnets. The outer portions of the garnets were sufficiently higher in 87Sr/86Sr compared to central rim portions that growth rates of the garnets (1.0 – 1.7 mm Myr21) and duration of total garnet growth (9 –13 Myr21) could be determined. Vance and O’Nions (1990) followed a similar procedure measuring both Rb – Sr and Sm –Nd isotopic parameters on single garnet sections obtained by sawing garnet crystals into inner, middle, and outer portions. The ,448 Ma isochrons that were obtained presumably established the timing of prograde growth of the garnets from that region of Newfoundland. In a different approach to determining the timing of prograde metamorphism, Burton and O’Nions (1991) analyzed the isotopic composition of small (1 mm) garnets from interlayered metasedimentary rocks that experienced a common P – T– t history, but one in which garnets formed at two different P – T conditions, as a function of different H2O activities. Based on the determined Rb – Sr isochrons the lower P– T garnets formed at 437.3 ^ 11.4 Ma and the higher P – T garnets at 423.5 ^ 4.7 Ma, in excellent agreement with the determined Sm – Nd isochrons (434.1 ^ 1.2 Ma and 424.6 ^ 1.2 Ma, respectively). These ages, when combined with the paleothermometric and barometric data, provided a significant amount of information on the rate of metamorphic processes in this region of Norway. These types of studies in similar regions of crustal thickening should provide an advance in our understanding of the lithospheric thermal response during collisional tectonics. With the demonstration that extremely small amounts of meteoritic material could be dated via Rb – Sr “microchrons” (Papanastassiou and Wasserburg, 1981) the path was opened for microanalysis of terrestrial samples, allowing much more control over the Rb – Sr dating of metamorphic material than previously possible. By analyzing small quantities of white mica from metamorphosed rocks the Rb – Sr system is capable of providing ages of formation, rather than cooling ages, thus establishing the time of at least one specific event in the metamorphic history of an area (e.g., Cliff, 1994; Chen et al., 1996). These types of studies have been further refined and the more recent microchron methods employ

Determination of Ages of Uplift or Exhumation a micro-drill and petrographic microscope to allow very specific areas of a geological thick section (,50 mm) to be sampled with complete textural control. Mu¨ller et al. (2000a) demonstrated the power of this technique by determining Rb – Sr dates of white mica from mylonites, which developed under greenschist facies metamorphic conditions, from shear zones in the eastern Alps. In an even more novel approach, the timing of the duration of shearing was established by determining Rb –Sr dates of micromilled samples of crystal fibers that developed in the strain fringe around pyrite grains from a fault zone in the northern Pyrenees (Mu¨ller et al., 2000b). These types of studies are just at their beginning stages but appear to be poised to revolutionize our ability to determine the timing, and possibly duration, of different deformational events, a critical step to more fully understanding all aspects of crustal evolutionary processes. 3.10.3.3 Sm–Nd and Lu–Hf Dating The main target of Sm –Nd and Lu –Hf dating in metamorphic studies continues to be the mineral garnet. Garnet is a major constituent in many metamorphic rocks; it preferentially incorporates heavy rare earth elements, and hence can have very high 147Sm/144Nd (e.g., Stosch and Lugmair, 1987) and 176Lu/177Hf ratios (e.g., Ducheˆne et al., 1997), and it has been widely used in thermobarometric studies, thus potentially providing a direct link between time and P –T conditions. In addition, garnet has a relatively high closure temperature for both the Sm –Nd (Mezger et al., 1992; Ganguly et al., 1998) and Lu – Hf systems (Scherer et al., 2000), thus increasing its attraction as a useful mineral in determining the timing of metamorphic events. Early studies were geared towards determining the timing of garnet growth during prograde metamorphism and relied on analyzing very large garnets (Vance and O’Nions, 1990; Mezger et al., 1992; Getty et al., 1993). Caution must be used in interpretation of these types of analyses, however, as it may be inclusions of REErich accessory minerals, and not the garnet itself, that dominates the Sm – Nd budget (see De Wolf et al., 1996 for a discussion). These types of analyses can provide information about garnet growth only if the inclusions and host garnet grew simultaneously or if the inclusions were isotopically equilibrated with the host matrix. In a more recent study, cores of up to 50 single garnets were mechanically isolated and then combined for Sm –Nd analysis in an effort to minimize the effects of averaging of different zones and thus provide better constraints on the timing of peak metamorphism (Argles et al., 1999). With the increased availability of computer-controlled microdrilling devices, ones capable of isolating

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very narrow regions of silicate minerals (e.g., Mu¨ller et al., 2000a), it is likely that future studies will focus on selecting even more specific regions within single garnets for thermochronology. The ability to select specific intracrystalline regions for analysis (e.g., Ducea et al., 2003), combined with increasingly sophisticated leaching techniques to minimize the effect of microinclusion contamination (e.g., Amato et al., 1999; Anczkiewicz et al., 2002), should increase the accuracy and precision of garnet thermochronology. Similarly, the increasing number of isotope laboratories with MC – ICP – MS instruments capable of analyzing very small quantities of hafnium will likely cause a dramatic increase in the number of metamorphic studies employing Lu –Hf garnet dating. When such studies are combined with the recent advances in experimental studies of REE diffusion and reexamination of concepts of closure ages (Ganguly et al., 1998; Ganguly and Tirone, 1999, 2001; Albare`de, 2003), substantial progress should be made in our understanding of the timing and duration of prograde versus retrograde metamorphic reactions. P – T –t studies will thus become ever more important to crustal evolution studies as a whole as the link between geochronology and metamorphic textural context becomes increasingly strengthened (e.g., Mu¨ller, 2003).

3.10.4 DETERMINATION OF AGES OF UPLIFT OR EXHUMATION Determining the magnitude and timing of crustal uplift or exhumation of orogenic belts are critical to our understanding of the crustal evolution of the regions investigated. Under favorable conditions, the magnitude of the exhumation of part of the crust can be estimated by geobarometry. If geobarometric information is combined with measured “cooling ages” of different minerals with very different closure temperatures, then the average rate of crustal exhumation can be estimated. Of most interest to the majority of uplift/exhumation studies are relatively low-temperature (50– 300 8C) thermochronologic techniques. Three such techniques, from highest to lowest closure temperature, are discussed.

3.10.4.1

40

Ar/39Ar Dating of Potassium Feldspar

Most common potassium-bearing minerals lose variable amounts of radiogenic argon at geologically modest temperatures, which at first glance would appear to make 40Ar/39Ar dating of limited geochronological use. However, argon diffusion appears to be a thermally activated process thus

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Ages and Growth of the Continental Crust from Radiogenic Isotopes 40

making the Ar/39Ar technique an excellent and widely used thermochronometer (see Section 3.10.3.1 and Chapter 3.08). One of the most important recent applications of 40Ar/39Ar dating to apparent uplift/exhumation studies is potassium feldspar thermochronology. An early assumption was that single closure temperatures (Tc) could be determined for potassium feldspars from slowly cooled plutons by using the 40Ar/39Ar data from the lower temperature (,600 8C) steps of a step-heating analysis (e.g., Heizler et al., 1988), consistent with single diffusion domains within the feldspars. Lovera et al. (1989) demonstrated inconsistencies with cooling rates determined by the specific closure temperatures and those calculated by examining the release spectra and thus suggested that the feldspar diffusion domains were of variable size, a suggestion also made by Zeitler (1988). In subsequent work, Lovera and Richter (1991) further demonstrated the multidomain behavior offeldspars by demonstrating the same phenomenon even when analyzing single crystals. Of particular importance was the demonstration that modeling the thermal history of feldspar was largely independent of domain geometry, size, and volume fraction. Thus, by measuring the 39Ar release spectra from a large number of heating steps of potassium feldspar from a single rock, and generating Arrhenius plots from those data, an apparently very robustly modeled cooling history can be obtained. A wide number of such cooling studies have now been made (see McDougall and Harrison (1999) for full references). Such studies when used in conjunction with other low-temperature techniques such as apatite fission track (FT) analysis and (U–Th)/He dating (see below) can provide a very substantial percentage of the full cooling path encountered by specific regions of crust. 3.10.4.2 FT Dating of Apatite In 1962 it was demonstrated that by chemically etching a uranium-bearing mineral the paths, or tracks, traveled by the fragments arising from the spontaneous fission of 238U could easily be viewed with an optical microscope (Price and Walker, 1962). A year later the first FT date of a mineral had been published (Price and Walker, 1963). The track, caused by disruption of the crystal lattice from the oppositely moving fission fragments, is initially ,16– 17 mm in length in apatite crystals (after etching), but becomes increasingly shortened, or annealed, with both increasing time and temperature (e.g., Gleadow et al., 1986). The maximum temperatures that can be reached and still retain abundant FTs vary with different minerals, in the same way as does the closure temperature for retention of a specific radiogenic daughter isotope. For apatite, such a closure temperature with regard to FTs is generally quoted

at ,100 8C (e.g., 105 ^ 10 8C; Parrish, 1983). Assigning an FT closure temperature is much less straightforward than the retention of a daughter isotope, however, as there are no geological temperatures, even 20 8C, at which annealing can be considered negligible (Donelick et al., 1990). Also, chemically different apatites can exhibit significantly different annealing behaviors (e.g., Green et al., 1989). Because FT formation can be viewed as a continuous, constant process each track has the potential to experience a different segment of the thermal history of the rock containing the apatite crystal. Thus each track could become shortened (annealed) to varying degrees, depending on the thermal history, and if accurate models of FT annealing can be made then the distribution of track lengths can provide considerable insight into that thermal history (e.g., Green et al., 1986, and references therein). A primary objective of annealing experiments is therefore to establish a thermal model that describes the behavior of FT systematics over geological time scales (e.g., Laslett et al., 1987, and many others). Fitting FT data from natural samples to such models may allow significant constraints to be placed on the past uplift/exhumation history of a crustal region, one of the most significant goals of FT thermochronology. Such studies have been applied to the unroofing history of a large number of orogenic belts (e.g., Fitzgerald et al., 1995; Gallagher et al., 1998). A potential problem with some of the earlier annealing models, upon which most FT thermal studies of sedimentary basins have been based, is that they characterize only a single type of apatite (i.e., the models are monokinetic), which may not always be applicable given the demonstration of different annealing properties of apatites of different composition (Green et al., 1989; Carlson et al., 1999). Significant advances have being made in multikinetic thermal modeling of apatite FT annealing (Ketcham et al., 1999), which has the potential to significantly advance the level of modeling of sedimentary basin thermal evolution as well as refine further crustal uplift and exhumation studies. 3.10.4.3 (U–Th)/He Dating of Apatite Some of the earliest attempts at dating uraniumbearing minerals were made by measuring the accumulation of helium in crystals from the a-decay of uranium and thorium. Two efforts that immediately predated the development of modern geochronology were Hurley (1954) and Damon (1957). Because of the ease of helium diffusion, however, the dates calculated were shown to be too young in most cases and the technique was soon abandoned in favor of

Neodymium Isotopes and Chemical Age of Crust U –Th – Pb isotopic techniques. A resurgence of interest in the technique began in the late 1980s with a more thorough quantitative understanding of the diffusive behavior of helium in different minerals (Zeitler et al., 1987). It has now been demonstrated that the mineral apatite has a closure temperature for helium of ,70 8C (Wolf et al., 1996) and thus is well suited as a very lowtemperature thermochronometer, that can further extend information relating to a variety of uplift and shallow crustal studies (see reviews by Farley, 2002; Ehlers and Farley, 2003). The typical method of determining a (U –Th)/He date is by extraction of helium gas either with a furnace (e.g., Zeitler et al., 1987; Lippolt et al., 1994; Wolf et al., 1996) or by laser (e.g., House et al., 2000) followed by mass spectrometric analysis, usually with small quadrupole-based instruments. After helium extraction, the apatite grains are recovered and dissolved for measurement of uranium and thorium abundances. There does not appear to be any loss of thorium or uranium during the vacuum extraction of the helium. Before the timing of helium closure can be calculated from the collected data, a correction factor must be applied because of the phenomenon of a ejection (see Farley et al., 1996). The a particles produced from uranium and thorium decay can travel 20 mm through an apatite crystal lattice, thus a particles will be ejected (i.e. helium loss) when the parent nucleus occurs near the edge of the crystal. Corrections for this helium loss must be estimated, and are based on the assumptions of idealized crystal geometry and near homogeneous distribution of uranium and thorium. The correction factors to the age, based on the size of the crystal, are typically between 1.2 and 1.5 (Ehlers and Farley, 2003), with uncertainties in the correction factors increasing with decreasing crystal size. For moderate to large apatite crystals (i.e., those occurring in typical plutonic rocks) the reproducibility of helium cooling ages, combining analytical errors with uncertainties in a ejection correction factors, is ,^5% (Ehlers and Farley, 2003).

3.10.5 NEODYMIUM ISOTOPES AND CHEMICAL AGE OF CRUST 3.10.5.1 Sm–Nd Methodology The Sm – Nd isotopic method depends upon the decay of 147Sm, comprising ,15% of natural samarium, to 143Nd by a-decay. With a reasonably well-known decay constant of 6.54 £ 10212 yr21 (Lugmair and Marti, 1978; Begemann et al., 2001), production of 143Nd is slow. The ratio 143Nd/ 144Nd, which is measured by isotope geologists, changed from ,0.50687 at the birth

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of the Earth to ,0.51264 today (Jacobsen and Wasserburg, 1984), with the highest values of major rock reservoirs under mid-ocean ridges possessing values ,0.5132. Because the precision of measurement on this ratio is typically ,5 £ 1026 from modern mass spectrometers, subtle variations of 143Nd/144Nd are actually quite easily discernible. Samarium and neodymium are both REEs. They therefore belong to a series of elements that have a very important role in geochemistry, due to progressive chemical fractionations that occur between the lighter and heavier REE. Although samarium and neodymium are adjacent elements in the naturally occurring REE series, fractionation between them is a little larger than for two elements with sequential atomic numbers, because there is a missing element between them, promethium, which has no stable isotope. The utility of the Sm – Nd isotopic method in crustal history is driven by the fact that upper continental crust acquires, due to igneous differentiation, a parent/daughter ratio 147Sm/144Nd that is ,45% lower than that of undifferentiated Earth, and lower still compared to typical depleted upper mantle sources. This fractionation is due to mantle minerals such as garnet, clinopyroxene, and orthopyroxene having lower distribution coefficients for light REE than for heavy REE, so that neodymium is partitioned into magmas slightly more strongly than samarium when mantle sources are melted. Because the 147Sm/144Nd ratio tends to be reasonably constant in average upper crustal rocks like granite, felsic gneiss and shale, and always ,45% lower than upper mantle values, the evolution of 143Nd/144Nd in upper crustal rocks slows down compared to the undifferentiated Earth and to upper mantle reservoirs, always by about the same amount. Thus, the 143Nd/144Nd ratio measured in a crustal rock is usually a good reflection of the average age of mantle separation of the materials in the rock. The rather constant fractionation of Sm/Nd ratios in upper continental crustal rock reservoirs is the basis for the widely applied neodymium model age that is illustrated in Figure 3. The Sm – Nd systematics of chondritic meteorites serve as a reference for the parent/daughter ratio of the undifferentiated Earth (Jacobsen and Wasserburg, 1984), labeled as CHUR for “chondritic uniform reservoir.” The evolution of this undifferentiated Earth is the basis for calculation of CHUR model ages (McCulloch and Wasserburg, 1978), while the neodymium isotopic evolution of the depleted upper part of the mantle is a more valid reference for most crustal materials, resulting in the DM model age (DePaolo, 1981). Neodymium isotopic compositions are usually given by 1Nd, where the deviations of 143Nd/144Nd above or below

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Figure 3 Sm – Nd systematics for crustal model ages and 1Nd values. The parameter 1Nd ¼ 104 :½143 Nd=144 NdðsampleÞ 2143 Nd=144 NdðCHURÞ=143 Nd=144 NdðCHURÞ; where all 143Nd/144Nd values are specified at the age of interest (t). CHUR Nd isotopic values can be obtained from the following equation: 143 Nd/144Nd(CHUR,t) ¼ 0.512638-0.1966.(elt-1), where l ¼ 6.54 £ 10212 a21. All crustal materials evolve towards negative 1Nd with time, because the evolution of 143Nd/144Nd is slower than the chondritic or bulk silicate Earth reference (CHUR). Nd model ages of crust can be calculated based on the chondritic reference (TCHUR) or on the intersection with the approximate evolution of depleted upper mantle sources (TDM). TDM is generally more meaningful for juvenile crustal rocks produced during plate-tectonic cycles, because it models the average separation age from the type of mantle that is commonly observed as a source of both oceanic crust and island arcs today. Note that the depleted mantle evolution is here shown as a straight line, but is very slightly concave-upwards in the model of DePaolo (1981).

CHUR are given as parts in 10,000 (see Figure 3 and caption). Because of their Sm/Nd ratios below CHUR (see Chapter 3.01), all continental crustal reservoirs evolve towards negative 1Nd with time. It is important to understand that whether neodymium model ages are an explicit part of the discussion (as in DePaolo, 1981), or are deemphasized in favor of an interpretation based on 1Nd values (e.g., Patchett et al., 1999), the utility of the Sm – Nd method in crustal evolution is ultimately based on the reproducible Sm/Nd fractionation that occurs between crust and mantle. Note that the Sm/Nd fractionation of more primitive crust of intermediate composition (Figure 3), such as basaltic andesite or andesite, does not evolve to negative 1Nd so rapidly as fully differentiated upper continental felsic crust. In this way, the neodymium isotopic composition of a crustal rock depends on the average mantle differentiation age of the components that went into making it. The origin of some crustal rocks can be very complicated, e.g., a granitoid melted from complex lower crustal sources (Chapter 3.11), or a sedimentary rock derived from erosion of multiple terrains. Thus, a single model age for a granitoid or a sedimentary rock can be quite misleading in terms of real geological events and processes. Nevertheless, it is the differences in age of separation from mantle

sources that bestow usefulness on neodymium isotopic variations, and no matter how complex the origin of a crustal rock, it is those age differences that are the ultimate basis for any interpretation. This is clearly seen, e.g., in the similarity of Sm – Nd systematics in many of the world’s major rivers today (Goldstein et al., 1984). Lu –Hf isotope systematics provide an important complement to Sm – Nd in the study of the crust and mantle (e.g., Patchett et al., 1981; Salters and Hart, 1991; Vervoort and Blichert-Toft, 1999). In the crustal context, Lu –Hf is extremely important because of the ,1% hafnium content of zircon, and the consequent ability to isotopically characterize the hafnium within grains that have been U – Pb dated (Patchett et al., 1981; Corfu and Stott, 1996; Vervoort et al., 1996; Amelin et al., 1999). However, the Lu – Hf isotopic system is currently overshadowed by a controversy over the decay constant. For many years, a value for the 176 Lu decay constant of 1.94 £ 10211 yr21, based on the eucrite meteorite isochron of Patchett and Tatsumoto (1980) and Tatsumoto et al. (1981) was used. More recent physical determinations reviewed by Begemann et al. (2001) have high dispersion, but do not seem to corroborate the 1.94 £ 10211 value. At the present time, there is a discrepancy between values based on U –Pb-dated terrestrial Precambrian REE-rich minerals, such

Neodymium Isotopes and Chemical Age of Crust as apatite, which suggest a decay constant of 1.865 £ 10211 (Scherer et al., 2001), and meteorite isochrons, that suggest values of (1.93 – 1.98) £ 10211 (Bizzarro et al., 2003; Blichert-Toft et al., 2002). For this reason, we mostly do not include Lu –Hf isotopic data in discussion of crustal age and origins. Lu –Hf data are of considerable importance in studies of early Archean rocks (see Section 3.10.6.4), and this uncertainty should be resolved as rapidly as possible.

3.10.5.2 Juvenile Crust Production versus Intracrustal Recycling It is fundamental to the neodymium isotopic approach that neodymium isotopes are able to distinguish between material added newly to the Earth’s continents and material that is merely recycled older crustal rock. This is important because sedimentary rocks derived from erosion of older continent may appear quite similar to those derived from erosion of young island-arc terrain, and unless they show marked S-type characteristics, granitoids are notoriously similar to each other, regardless of their ultimate origin. Thus, in a world where all regions of continents had been characterized for both orogenic ages and neodymium isotopic characteristics, one could draw two global maps. One would show orogenicbelt ages, representing times of consolidation of regions of the continents, while the other would show generally older neodymium-based average ages, that would represent the true differentiation age of the crust. Some orogenic belts might consist of dominantly juvenile crust, with

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neodymium ages similar to the orogenic age, but other belts might be dominated by materials recycled from older crustal terrains. Because the sedimentary system is a powerful mover of crustal detritus over large geographic scales, and because melting to produce granitoid batholiths often averages large domains of lower continental crust, orogenic belts very often have mixed origins in terms of the crustal age of their components. Regional maps of crustal age based on neodymium model ages, even without the references to the isotopic work, appear quite often in the literature and in presentations (e.g., Karlstrom et al., 1999). The coverage of U –Pb ages and neodymium isotopic data is not yet sufficient to draw robust global maps, and because the current pace of data accumulation is high, we do not attempt to compile global maps in this chapter. Instead, the approaches will be illustrated with examples.

3.10.5.3 Juvenile Crust Production at 1.9–1.7 Ga The abundant crust that was produced in the 1.9 –1.7 Ga interval has been the subject of numerous studies, and the evolution of that work illustrates important elements in study of the continental crust. Following the demonstration of a juvenile origin for the ,1.8 Ga crustal assemblage in Colorado (DePaolo, 1981), there followed a period in which neodymium isotopic data were gathered for numerous terrains in the northern continents (Figure 4). In North America, studies by Nelson and DePaolo (1984, 1985), Bennett and DePaolo (1987),

Figure 4 Regions of crust of different ages described in this chapter. Dot pattern represents 1.9– 1.7 Ga juvenile crust of North America, Greenland, and Europe. Diagonal lines represent juvenile crust of the Altaid collage of orogenic belts and of the Canadian– Alaskan Cordillera. Black rectangles are locations of small regions of pre-3 Ga crust: WG—west Greenland; La—Labrador; MN—Minnesota; Ac—Acasta Gneiss Complex, northwestern Canada; Swz—Swaziland; Nar—Narryer terrane, western Australia.

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Chauvel et al. (1987), Barovich et al. (1989), and Bowring and Podosek (1989) added rapidly to the database. In Greenland and Europe, work by Patchett and Bridgwater (1984), Skio¨ld and Cliff (1984), Kalsbeek and Taylor (1985), Patchett and Kouvo (1986), Huhma (1986), Kalsbeek et al. (1987), Claesson (1987) and Patchett et al. (1987) performed the same function. All of the studies cited above documented a high proportion of juvenile crust production in the 1.9 – 1.7 Ga orogenic belts of the northern continents, as highlighted in the reviews by Patchett and Arndt (1986) and Condie (1990). Large volumes of new crust were added, at least to the studied regions, during 1.9– 1.7 Ga. It appears that the 1.9 – 1.7 Ga crust of the northern continents is approximately equal in mass to what could be produced by present-day island-arc generation rates over 200 Myr (Reymer and Schubert, 1984; Patchett and Arndt, 1986). Patchett and Chase (2002) modeled this in terms of accumulation of juvenile terrains in restricted regions of the globe due to transform motions associated with large-scale plate tectonics. Since the late 1980s, two important changes have occurred in the studies of the growth of 1.9– 1.7 Ga continental crust. The first is that neodymium isotopic data have become a very common component of any tectonic study of orogenic belts, granitoid complexes or sedimentary sequences, particularly in the Precambrian, meaning that neodymium isotopic data are now routinely gathered as part of characterization of regions not studied before. The second change is that the pace of data acquisition has slowed a little, because regions that were well-characterized in terms of geology and geochronology were studied in the 1980s, while fresh mapping and geochronology are needed to open up new areas for credible neodymium isotopic study. Thus, more recent studies incorporating neodymium isotopes are not super-regional reconnaissance studies of regions for which geology and ages are already known, but are targeted to newer topical problems and areas in Proterozoic geology, such as ophiolites in Quebec and Finland (Hegner and Bevier, 1991; Peltonen et al., 1998), collisional ¨ hlander et al., 1999), or granitoid tectonics (O petrogenesis (Valbracht et al., 1994; Krogstad and Walker, 1996; Ra¨mo¨ et al., 2001). In regions away from North America, Europe and Australia, the pace at which the volume both of well-dated Precambrian terrain and of neodymium isotopic information increases is often much lower. This leaves large parts of the continents poorly known for neodymium crustal age. The very important conclusion, that the 1.9– 1.7 Ga period saw a sequence of relatively rapid and high-volume additions to the continental crust, is not likely to change. However, it is not

yet really clear if the 1.9 –1.7 Ga crust of the northern continents could represent juvenile terrain gathered from the rest of the globe, as suggested by Patchett and Arndt (1986) and Patchett and Chase (2002). That would be a model where the 1.9 –1.7 Ga period could have shown normal crustal growth rates, but where plate tectonics had grouped the juvenile crustal products into restricted regions. If the continental regions that are currently poorly known for their crustal age also show significant juvenile crust of 1.9 – 1.7 Ga age, then this event must have been larger than presently documented in the northern continents, and would represent a distinct, positive spike in crustal growth. Such a situation might correspond to crustal growth initiated by mantle plume activity, as hypothesized by authors like Boher et al. (1992) and Stein and Hofmann (1994).

3.10.5.4 Juvenile Crust Production in the Canadian Cordillera The Canadian –Alaskan Cordillera (Figures 4 and 5) has become the typical example of a Phanerozoic orogenic belt that contains significant amounts of juvenile crust. It is distinctive because of all the global history of mountain-building events from Cambrian to Cretaceous time, only three orogenic zones seem to display significant juvenile additions. Two of these are the South Island of New Zealand (e.g., Frost and Coombs, 1989), which is small in size, and the very large Altaid orogenic collage of central Asia (e.g., Sengo¨r et al., 1993), for which isotopic data have only recently become available. The Canadian – Alaskan Cordillera has been studied for radiogenic isotopes by several groups against a backdrop of rather well-constrained geology. In the 1970s, the Canadian Cordillera became a center of tectonic attention because of the demonstration of the distinctness of the belts of rocks from which it was constructed (Monger et al., 1972). This region became the typical example for the concept of tectonostratigraphic terranes and their assembly by accretion to form a mountain belt (Coney et al., 1980). Subsequently, discussions were initiated about the apparent transpressional emplacement of many of the terranes, and northward transform faulting of them after emplacement (Gabrielse, 1985; Umhoefer, 1987). A recent geological review of Cordilleran geology and the development of ideas are found in Monger (1993). In parallel with the geological investigations, geochronology and strontium isotopic tracer investigations were carried out by Armstrong and co-workers (Armstrong, 1988; Armstrong and Ward, 1993). Because initial 87Sr/86Sr ratios of

Neodymium Isotopes and Chemical Age of Crust

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Figure 5 Terranes of the Canadian Cordillera. Older continental components are represented by the Miogeocline, which rests on North American Precambrian basement, and the Yukon – Tanana terrane. Of the other terranes, those near to the Miogeocline, such as Kootenay, have major older crustal components. Juvenile characteristics increase in the direction of the Pacific Ocean (after Patchett et al., 1998; Butler et al., 2001).

plutonic rocks in the outboard parts of the Canadian Cordillera tended to be low, and because geological associations suggested that a high proportion of intra-oceanic rock assemblages had been accreted to North America, there was a general supposition that juvenile crustal elements from the paleo-Pacific basin would be abundant in the Canadian Cordillera (Monger et al., 1972; Coney et al., 1980; Armstrong, 1988). However, many of the terranes, even outboard ones, contain Paleozoic rock assemblages that were viewed as “continental” (Monger et al., 1972). It was against this background that the present authors, in collaboration with George Gehrels, initiated neodymium isotopic characterization of the Canadian Cordillera, beginning with the

Alexander terrane (Figure 5). In this outboard terrane, tightly folded Neoproterozoic and Paleozoic felsic igneous assemblages form the older components, and were some of the rocks referred to by Monger et al. (1972) as “continental.” At the time when the first research proposal was written, it seemed to the present first author that these assemblages would yield neodymium isotopic data typical of older Precambrian continental crust. The outcome was very different, because all components of the Alexander terrane, including the oldest ones, turned out to have positive initial 1Nd values (Samson et al., 1989). This result became the keynote for much of subsequent isotopic study by our group (Samson et al., 1990, 1991a,b). The isotopic work was consistent with outboard terranes being generally

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more juvenile, but in addition it was consistent with ancient sedimentary assemblages having been incorporated into the Coast Plutonic Complex (Samson et al., 1991a,b; Boghossian and Gehrels, 2000; Patchett et al., 1998). By no means did these authors have a monopoly of isotopic study in the Canadian – Alaskan Cordillera, and important data for outboard terranes, the Coast Plutonic Complex (Figure 5) and related areas were published by Arth et al. (1988), Farmer et al. (1993), Mezger et al. (2001), Cui and Russell (1995), Friedman et al. (1995) and Mahoney et al. (1995). One important finding of the last three cited papers is to show that the southern part of the Coast Plutonic Complex does not contain the older crustal elements seen further north around the Alaska Panhandle. Further inboard lies a zone of the Canadian Cordillera that incorporates terranes showing well-known oceanic features, like Cache Creek and Slide Mountain, arc assemblages like Quesnel, and also large amounts of clastic sedimentary rocks, along with the Yukon– Tanana terrane of more continental affinity in the northern part of the Cordillera (Figure 5). This zone is characterized by juvenile neodymium signatures in most of its volcanic rocks and mafic intrusives (Smith and Lambert, 1995; Smith et al., 1995; Piercey et al., 2002). Neodymium isotopic signatures showing a mixture of juvenile and Precambrian continental North American materials are shown by both clastic sedimentary assemblages (Creaser et al., 1997; Patchett and Gehrels, 1998; Unterschutz et al., 2002; Erdmer et al., 2002), and granitoid plutons (Ghosh, 1995; Ghosh and Lambert, 1995; Piercey et al., 2003). Still further inboard towards North America lies the Omineca Crystalline Belt (Monger et al., 1982) and the Miogeocline, where pre-Cordilleran Neoproterozoic through Late Jurassic sedimentary rocks rest on North America basement. The Omineca crystalline belt is not separately distinguished on Figure 5, because it is a metamorphic/ plutonic overprint rather than a separate terrane, but it occupies the general area of the Quesnel and Kootenay terranes. Although juvenile neodymium is sometimes seen, the Omineca belt generally shows neodymium isotopic signatures, in both metasedimentary and plutonic rocks, that correspond to older North American continent (Burwash et al., 1988; Ghosh and Lambert, 1989; Stevens et al., 1996; Brandon and Lambert, 1993, 1994; Brandon and Smith, 1994; Driver et al., 2000). The miogeoclinal sequence appears to have been derived from Precambrian basement before ,450 Ma, but after this time to have been supplied by distant mountains of early- to mid-Paleozoic age in the Canadian Arctic (Boghossian et al., 1996; Garzione et al., 1997; Patchett et al., 1999).

The general picture of the Canadian – Alaskan Cordillera is of a mountain belt characterized by accretion to older continental crust of juvenile crustal elements from the ocean floor, or from island-arc environments, whose proportion increases towards the Pacific Ocean. Generally, work in the Canadian –Alaskan Cordillera has evolved in the same way as work on Proterozoic terrains described above. From initial reconnaissance studies and large-scale views of how the crust grew, research has evolved into more detailed studies with more precise geological control. Nevertheless, some important general issues about crustal evolution come to fore by thinking about the Cordillera at the largest scale. Samson and Patchett (1991) reviewed the thenexisting neodymium isotopic database, and concluded that ,50% by mass of the Canadian segment of the Cordillera was juvenile crustal material. Clearly, such estimates are dependent on assumptions about whether terranes continue to lower crustal depth. Seismic data often reveal that juvenile belts are underlain by older continental material, both in Proterozoic and Phanerozoic contexts (Lucas et al., 1993; Clowes et al., 1999). The 50%-juvenile estimate of Samson and Patchett (1991) should certainly be revised downwards in light of seismic results summarized by Clowes et al. (1999), as well as more recent field and isotopic work such as that of Erdmer et al. (2002). Qualitatively however, the status of the Canadian Cordillera as the locus of juvenile crustal growth remains in spite of these uncertainties. Samson and Patchett (1991) also viewed the Canadian Cordillera as an analogue for Proterozoic crustal growth by accretion of juvenile terranes, a comparison that was explicitly examined in terms of major- and trace-element compositions by Condie and Chomiak (1996). The present authors, like many others, have always been impressed by the transform faulting that is presently active along the North American margin, which serves to pile juvenile crust into one region that stretches from southern British Columbia to Alaska. Subsequent synthesis of Cordilleran geology (Johnston, 2001) and discussions about crustal growth in general (Patchett and Chase, 2002) have emphasized the northward along-margin transport displayed in the Canadian Cordillera as a critical element in crustal growth models. The rationale for this is that along-margin transport of slices of juvenile crustal material is able to pile them all up in one segment of the continental margin, perhaps even into a restricted region of the globe, creating the impression of a very intense and localized crustal growth period (Patchett and Chase, 2002), for which extraordinary mechanisms might otherwise be required (e.g., Stein and Hofmann, 1994). This process

Isotopes and Pre-3 GA Continental Crust should have been important in most of geologic time, accounting for numerous apparently very intense crustal growth episodes (Patchett and Chase, 2002). 3.10.5.5 Juvenile Crust Production in the Altaid Collage of Central Asia Another region of Phanerozoic crustal growth is the Altaid orogenic collage of central Asia, located in the eastern parts of the former Soviet Union, and in western China (Figure 4). The rocks range in age dominantly from Neoproterozoic through Early Carboniferous, with final tectonic movements and post-orogenic igneous activity extending through part of Permian time. Thus, the crustal assemblage is older than the bulk of the material in the Canadian Cordillera. In a simplified view, one might say that the Altaid system represents the bulk of global juvenile crustal accretion of Cambrian through Permian time, while the Canadian Cordillera and New Zealand have that role from Triassic through Cretaceous time. Early work on the tectonic assembly of Asia, which was published in international journals, took place in the 1980s and early 1990s (e.g., Sengo¨r and Hsu¨, 1984; Windley et al., 1990; Allen et al., 1992). However, the Altaid system came to the attention of the wider scientific community through the bold syntheses of Sengo¨r and coworkers (Sengo¨r et al., 1993; Sengo¨r and Natalin, 1996a,b). These authors described a collage of belts of arc volcanic rocks, ophiolitic assemblages, greywacke-shale sequences and granitoids, that occupies a very large area in central Asia to the north of the Himalayan system, and lying between the Baltic Shield/Russian Platform to the west, the Siberian Shield to the east, and with the North China Craton providing a partial boundary to the south. Tectonic trends are highly variable, and this seems consistent with the interpretations of Sengo¨r and co-workers, that the collage represents the sweepings of a large ocean, with telescoping of accreted fragments not only by orthogonal collision, but with transpression and transform faulting, essentially along the strike of the developing orogen, over very large geographic scales. This resulted in a collage with variable tectonic trends occupying a large region of roughly triangular shape in the center of the Asian continent. Isotopic analyses of Altaid-collage rocks are not as abundant as those in the Canadian Cordillera, but have appeared in increasing numbers since about 1996. Kovalenko et al. (1999, 2003) presented neodymium isotopic data on juvenile Altaid rocks in Russia and Mongolia, while Wu et al. (2000), Jahn et al. (2000) and Chen and Jahn (2002) made extensive studies of

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the abundant granitoid plutons in the Chinese part of the Altaid assemblage. Yarmolyuk et al. (1999) and Hu et al. (2000) used neodymium isotopic results to study the nature of pre-Altaid basement. Modern U – Pb zircon geochronology, which together with detailed modern documentation of tectonic relations in the field, has also been largely lacking in the Altaid crustal collage, is being actively undertaken (Wilde et al., 2000; Salnikova et al., 2001; Windley et al., 2002). All authors concluded that substantial juvenile material was added to Asia during Altaid events, particularly during Early Ordovician through Devonian time. Post-orogenic granitoids of Late Carboniferous to Early Permian age have sources that may appear juvenile, or to consist of older continental crust, depending on the location with respect to the margins of the Altaid collage (Han et al., 1997; Litvinovsky et al., 2002). What appears to emerge from the ongoing studies is that the Altaid collage consists of terranes of juvenile material juxtaposed with terranes containing pre-existing Proterozoic crustal rocks over the whole extent of the belt. Sengo¨r et al. (1993) painted just such a picture, and also suggested that up to 50% of the entire Altaid collage may represent juvenile crust of Paleozoic age. So far, the neodymium isotopic data appear to support juvenile crust of about this magnitude, but much field data, geochronology and neodymium isotopic data remain to be gathered. In addition, estimates for the volume of juvenile crustal material may need revision when better data are available for whether and how the juvenile terranes project into the lower continental crust, as noted above in the context of the Canadian Cordillera. Estimates of the proportion of juvenile crustal growth represented by the Altaid collage can be expected to be refined in the future.

3.10.6 ISOTOPES AND PRE-3 Ga CONTINENTAL CRUST 3.10.6.1 Existence of Ancient Continental Crust Isotope geology acquires a dominant significance when continental rocks dating from before 3 Ga are studied. This is because fossils are not available, and metamorphism and tectonism are often so severe that original layering of supracrustal sequences may not be visible, and even magmatic contact relationships may be obscured. Gneisses of 3.8 Ga, 2.8 Ga, and 1.8 Ga may look very similar in the field, and isotopic dating is absolutely required to distinguish them. In addition, pre-3 Ga rocks may carry unique information about environments on the early Earth, or about mantle differentiation and layering, and the search for the oldest continental crust has always

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been a major area of activity. All the areas of pre3 Ga crust discussed here are indicated on Figure 4. Early dates for ancient crust were obtained by the Rb – Sr whole-rock isochron method, or by the U – Pb zircon method when it was in its infancy. Approximately 3.5 Ga U –Pb zircon ages were obtained from gneisses in Minnesota by Goldich et al. (1970), and gneisses in western Greenland were dated at 3.6 –3.8 Ga by Rb – Sr isochrons (Black et al., 1971; Moorbath et al., 1972). These age determinations were the major scientific news of their day in geology, and began a whole field of endeavor in deciphering the history of ancient crustal rocks. Early foci of research on ancient crust were western Greenland (Moorbath, 1978), Labrador (Hurst et al., 1975; Bridgwater and Collerson, 1976) and Swaziland (Davies and Allsopp, 1976). Early studies of these regions using Sm – Nd were by Hamilton et al. (1983) and Carlson et al. (1983), and these and other isotopic studies documented the juvenile nature of the these earliest crustal gneiss complexes. Presently, most occurrences of pre-3 Ga rocks have been accurately dated using modern approaches of the U – Pb zircon method (e.g., Kro¨ner et al., 1989; Bowring et al., 1989b; Horstwood et al., 1999), and neodymium isotope data have been widely obtained (e.g., Collerson et al., 1991). Hafnium isotopic data, combined with neodymium, have been obtained for critical samples (Vervoort et al., 1996; Vervoort and Blichert-Toft, 1999; Blichert-Toft et al., 1999; Amelin et al., 1999). The focus of neodymium and hafnium isotopic studies has moved away from documentation of juvenile character towards an attempt to constrain the degree of heterogeneity in the early Earth’s mantle, with a view to detecting “primordial” rock reservoirs and effects of the primary differentiation of the Earth (e.g., Bowring and Housh, 1995; Albare`de et al., 2000; see Chapter 2.13). Many arguments have developed concerning the reliability of isotopic values for single samples of polymetamorphic gneisses (e.g., Vervoort et al., 1996; Gruau et al., 1996; Moorbath et al., 1997). The case of the Acasta gneisses, to which the above-cited references apply, is detailed later in this chapter.

3.10.6.2 Crustal Growth Events and Recycling into the Mantle Early in the evolution of isotopic studies of ancient crust, Moorbath (1975, 1978) developed the concepts of (i) major Precambrian crustal growth events in which juvenile crust was made, and (ii) the essential indestructibility of continental crust once consolidated by orogenic events.

Models based on neodymium and strontium isotopes that grew the continents through time were made possible by the arrival of the meteoritically constrained Sm – Nd system (O’Nions et al., 1979; Jacobsen and Wasserburg, 1979; DePaolo, 1979). A very readable account of the chemical relationships between crust and mantle is by Hofmann (1988), while one of the more mathematically complete treatments is by Alle`gre et al. (1983a,b). Although a neodymiumbased cumulative age curve for global continental crust cannot yet be drawn, the estimates for North America and Europe (Patchett and Arndt, 1986; Condie, 1990) may approximate the global picture. Alternatively, the modeling by Armstrong (1968, 1981) and Zartman and Doe (1981), initially strongly based on lead isotopes, advocated the position that during crustal growth events or during plate movement in general, continent could also be destroyed by the subduction of sedimentary rocks resting on the ocean floor. It was realized by the community that all crust and mantle geochemistry and isotopes were consistent with large amounts of crustal recycling, provided that more or less complete mixing occurred during mantle convection. Much of the discussion concerns the present-day Earth. However, the period before 3 Ga is critical to the discussion of crustal recycling because it is generally agreed that large amounts (perhaps 40 – 50%) of presently surviving crust came into existence during the immediately following period, between 3.0 Ga and 2.6 Ga (e.g., Condie, 1990). Either the pre-3 Ga Earth had much less felsic crust than existed after 2.6 Ga, or large amounts of pre-3 Ga crust were destroyed before or during the time that the 3.0 –2.7 Ga crust was produced. If it could be shown that massive amounts of pre-3 Ga crust were destroyed, then essentially all objections to large-scale continent recycling in the mantle would disappear. Attempts to test the effects of sediment subduction (Patchett et al., 1984), or whether pre-3 Ga crust was available to be recycled into later crust (Stevenson and Patchett, 1990), founder on two possibilities. One is that in the early Earth, aided by a hotter mantle and continuing bombardment from asteroidal objects, it may have been possible to destroy continental fragments wholesale, so that little trace of them remained at the surface. Another is that large amounts of sediment subduction would result in a very well averaged geochemical signature entering the mantle, so that distinctive features of oceanic sediments, such as enrichment or depletion in zircon and its unradiogenic hafnium, could not be used to argue against the process (Plank and Langmuir, 1998). At the same time, budgets for the mass of rock entering subduction zones appear to suggest that

Isotopes and Pre-3 GA Continental Crust sedimentary material has to disappear into the mantle (von Huene and Scholl, 1993), an argument used forcefully by Armstrong (1981). Consequently, a discussion that was once framed in terms of “growth or no growth” (e.g., Stevenson and Patchett, 1990; Armstrong, 1991; McCulloch and Bennett, 1994) may now focus on how much continent recycling occurs, and the magnitude of recycling compared to the growth seen in places like the Canadian Cordillera. The magnitudes of these fluxes are very important for crust –mantle evolution, but are quite difficult to determine accurately, because of the global basis required, and the deep parts of subduction zones being difficult to image precisely (Reymer and Schubert, 1984; von Huene and Scholl, 1993; see Chapter 2.11). 3.10.6.3 Acasta Gneisses, Northwest Territories, Canada The Slave craton in northwestern Canada is an Archean granite – greenstone terrain that is bounded on the east and west by Paleoproterozoic orogenic belts (see geological reviews of Henderson (1985) and Padgham (1985)). Crustal regions within the western portion of Slave craton contain granitoids older than 2.8 Ga (see Bowring et al., 1990, and references therein). The antiquity (.3.5 Ga) of the westernmost portion of the Slave craton has been known since the mid-1980s (Bowring and Van Schmus, 1984), but in 1989 it was demonstrated, via ion microprobe U – Pb dating, that zircons within a tonalitic gneiss (sample BGXM) from the Acasta River area were as old as 3.96 Ga (Bowring et al., 1989a), making the gneiss the oldest known terrestrial rock. An identical age was determined for a second gneiss (sample SP-405) with a granitic composition (Bowring et al., 1990). Stern and Bleeker (1998) reported 207Pb – 206Pb dates as old as 4.02 Ga from a gneiss collected from within the same outcrop as that sampled by Bowring et al. (1989a). More recent U – Pb geochronological work on the gneisses in the Acasta River area (now called the Acasta Gneiss Complex) has extended the known extent of the oldest known surviving continental crustal region, as zircons from two metatonalites and a metagranodiorite have nearly concordant U – Pb dates of 4,002 ^ 4 Ma, 4,012 ^ 6 Ma, and 4,031 ^ 3 Ma, respectively (Bowring and Williams, 1999). The neodymium isotopic composition of Acasta gneiss BGXM was investigated by Bowring et al. (1989b). These authors reported a value of 1Nd(3.96 Ga) ¼ 21.7 for a whole-rock analysis and 1Nd(3.96 Ga) ¼ þ 0.7 for an amphibolitic-rich layer taken from a hand specimen of BGXM. An indistinguishable value, 1Nd(3.96 Ga) ¼ þ 0.8, was reported for the granitic gneiss SP-405

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(Bowring et al., 1990). The corresponding TCHUR ages of these three samples are 4.1 Ga, 3.92 Ga, and 3.85 Ga, respectively. The variable 1Nd values of the Acasta samples were interpreted by Bowring et al. (1990) as reflecting original protolith heterogeneity, rather than the effects of metamorphic disturbance. The negative 1Nd value and TCHUR age of 4.1 Ga of the tonalitic gneiss was interpreted as evidence that at least this portion of the Acasta Gneiss Complex must have been derived from, or interacted extensively with, a crustal reservoir considerably older than 3.96 Ga. The same argument cannot be easily made for the granite orthogneiss, however, as it has a TCHUR age that is 100 Myr younger than its crystallization age. Lead isotopic ratios were determined for HF-leaches of alkali feldspars from the granitic gneiss and of plagioclase from the tonalitic gneiss (Bowring et al., 1990). The 206Pb:207Pb: 208Pb:204Pb ratios of the least radiogenic leaches, presumed to be closest to the initial lead isotopic composition of the rocks, were 14.12, 15.08, 34.01 and 14.96, 15.29, 33.53, respectively. The 206 Pb/204Pb and 207Pb/204Pb ratios of the three plagioclase leaches lie on a linear array with a slope equivalent to an age of 3.60 ^ 0.28 Ga. Because this age overlaps the 3.62 Ga of the unzoned zircon rims from the tonalitic gneiss, the feldspar lead isotopic composition was interpreted as reflecting homogenization of whole-rock lead isotopes during a 3.6 Ga metamorphic event. The high 207Pb/204Pb ratios of the leached feldspars were taken as evidence that the gneisses were derived from a reservoir characterized by a high U/Pb ratio prior to 3.96 Ga, such as significantly older continental crust that had not lost uranium or thorium (Bowring et al., 1990). Implicit in this argument is that the lead isotopic compositions of the HF-leached feldspars accurately represent the lead isotopic composition of the protoliths to the gneisses, and hence accurately reflect the source materials of the protolith. Hafnium and U –Pb isotopic ratios were determined for single zircons from two gneisses reported to be Acasta gneisses (Amelin et al., 1999). However, no details of the sampling location or composition of the gneisses were given, except that they were from “granitic” and “amphibolitic” gneisses. Two zircons from the amphibolitic unit have 1Hf(t) values of 24.7 and 20.6, and the zircon from the granitic unit has a reported 1Hf(t) value of 21.8. However, the significance of these data to the 4.0 Ga Acasta gneisses are uncertain, as the zircons analyzed by Amelin et al. (1999) yielded 207Pb/206Pb dates of 3,548 Ma to 3,565 Ma, similar to the 3.6 Ga dates determined for some structureless equant zircons and the unzoned outer portion of some of the 4.0 Ga crystals analyzed by Bowring et al. (1989a, 1990),

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which they interpreted as dating the timing of a second generation of zircon growth. Because of the extreme complexity in the zoning of zircons from the Acasta gneisses, hafnium isotopic compositions of whole zircon crystals might reflect only a weighted average of different isotopic compositions. All in all, the complexity of the isotopic data and their interpretation in the Acasta Gneiss Complex illustrate the difficulties inherent in study of planetary evolution using polymetamorphic gneisses. This is the ultimate cause for the controversies over relaibility of radiogenic isotopic parameters derived from these rocks (e.g., Vervoort et al., 1996; Gruau et al., 1996; Moorbath et al., 1997). 3.10.6.4 Narryer Terrane, Western Australia The oldest known portion of the Archean Yilgarn craton is the Narryer terrane, occurring in the northwestern part of the craton. The Narryer terrane contains 3.7 Ga gneisses, 3.7 Ga anorthositic rocks, and 3.6 Ga granites (see Myers, 1995, and references therein). Two important belts of metasedimentary rocks, thought to have been deposited ,3 Ga, occur within the Narryer terrane: the Narryer and Jack Hills belts. The metasedimentary rocks within the Narryer belt include metaconglomerates and quartzites that have preserved cross-bedding despite being metamorphosed at upper amphibolite to granulite facies conditions (Froude et al., 1983). Detrital zircons extracted from one of these quartzites were the first pre – 4.0 Ga terrestrial minerals identified (Froude et al., 1983). In the original study, four zircons out of 102 crystals analyzed from one quartzite sample yielded 207Pb-206Pb dates of 4.11 – 4.18 Ga. The other detrital zircons from that sample and a second quartzite yielded 207Pb – 206Pb dates between 3.75– 3.3 Ga, with two zircons having nearly concordant 3.1 Ga dates. The Jack Hills belt, broadly similar to the Narryer belt but metamorphosed to upper greenschist facies, contains metabasalts, chert, and banded iron formation interleaved with clastic metasedimentary rocks (Compston and Pidgeon, 1986). Detrital zircons from a ,3.1 Ga chert pebble conglomerate were analyzed using ion microprobe U – Pb techniques as part of a continuing search for ancient zircons. Compston and Pidgeon (1986) found detrital zircons with 207Pb – 206Pb dates ranging between 4.28 – 4.0 Ga, significantly extending the sampling region of the oldest known minerals. Supporting geochronological evidence for the presence of 4 Ga detrital zircons from these areas was provided by both the zircon evaporation method (Kober et al., 1989) and the ID-TIMS method (Amelin, 1998). In a more recent ion microprobe study of detrital

zircons from Jack Hills, Mojzsis et al. (2001) confirmed the existence of 4.28 Ga zircons by identifying two grains giving concordant U –Pb dates of 4,279 ^ 5 Ma and 4,280 ^ 5 Ma. An even older detrital zircon, 4,404 ^ 8 Ma, has been discovered from the Jack Hills conglomerate (Wilde et al., 2001), making this crystal only ,150 Myr younger than the estimated time of Earth formation. In addition to geochronological studies of the Earth’s oldest zircons, geochemical and isotopic studies have also been performed (Kinny et al., 1991; Maas et al., 1992; Amelin et al., 1999; Mojzsis et al., 2001; Peck et al., 2001; Wilde et al., 2001). Maas et al. (1992) demonstrated that both the older and younger detrital zircon suites from the Jack Hills conglomerate have trace-element compositions and mineral inclusions consistent with nucleation from a felsic magma. Kinny et al. (1991) were the first to attempt to determine the hafnium isotopic composition of the zircons, but the large uncertainties in the ion microprobe measurements (^ (5 –7) 1Hf units) precluded a detailed discussion of the petrogenesis of the parent magmas. In more recent work, Amelin et al. (1999) measured 37 Jack Hills detrital zircons for U – Pb ages and hafnium isotopes. The advantage of measuring U –Pb and hafnium isotopes from the same dissolved zircon crystal is that the degree of discordance can be determined, which in turn is important in assessing the likelihood of the presence of a xenocrystic core and/or metamorphic rims on the zircon. Zircons that are either concordant or not too discordant are the least likely crystals to contain multiple domains. Similarly, zircons that have suffered the least lead loss are the least likely to have experienced open system behavior for lutetium or hafnium (see Samson et al., 2003 for discussion). Of the 10 zircons identified by Amelin et al. (1999) that are .3.8 Ga, two have 1Hf(t) values that are further than ^1 1Hf units from CHUR. The bulk of these ancient zircons are thus consistent with derivation from a chondritic-like source, which at this age could include very slightly to nondepleted mantle, or juvenile crust that had recently been extracted from such a mantle. Two crystals, with 207 Pb – 206Pb dates of 3.82 Ga and 3.97 Ga, have 1Hf(t) values of 22.2 ^ 0.6 and 22.7 ^ 0.4, respectively. These values are far enough below the CHUR line to suggest that these zircons crystallized from slightly evolved magmas, assuming that the zircons contain only singleage domains and that the ages and isotopic compositions are accurate. If those assumptions are correct, then we are presented with the intriguing possibility that an isotopically enriched reservoir existed as early as 3.97 Ga, perhaps similar to the one that gave rise to the 4.40 Ga Jack Hills zircon, in contrast to the general conclusions

References reached by Vervoort and Blichert-Toft (1999). However, all these hafnium isotope arguments are complicated by the present controversy over the decay constant of 176Lu, described in Section 3.10.5.1. The presence of very early evolved crustal sources is reinforced by oxygen isotope data for the ca. 4 Ga Jack Hills detrital zircons. Mojzsis et al. (2001) and Wilde et al. (2001) presented, independently, the first ion-microprobe oxygen isotope analyses of the .3.9 Ga detrital zircons. Wilde et al. (2001) determined d 18O values of þ5‰ and þ7.4‰ (mean of two analyses each) for two different regions within the 4.40 Ga crystal and Mojzsis et al. (2001) determined d 18O values of about þ6‰ to þ10‰ for 4.3– 3.9 Ga zircons. Peck et al. (2001) confirmed this general range of values, but individual spot measurements did not exceed d 18O of þ8.6‰. The importance of these values is that because there is a fractionation of ,1.5‰ between zircon and granitic magmas (e.g., Valley et al., 1994), the estimated d 18O values of the parental magmas to the Jack Hills zircons are in the vicinity of þ 8.5‰ to þ 9.5‰ (Peck et al., 2001). Such values are well beyond those of typical mantle (þ 5.5‰ to þ 6‰), and are most consistent with granitic magmas having been derived in part from materials that were once exposed to supracrustal conditions (i.e., lowtemperature weathering or diagenesis). This suggests the very early presence of some sort of hydrosphere, including possibly a 4.4 Ga ocean (e.g., Valley et al., 2002).

ACKNOWLEDGMENTS The authors are grateful to Clement Chase and Robert Butler for help with map figures, and to Roberta Rudnick and Mark Schmitz for their thorough reviews of this chapter. Samson acknowledges Syracuse University for granting a sabbatical leave, during which this chapter was written. Patchett was supported by NSF-EAR0003343, and Samson by NSF-EAR-0106853.

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subducted at convergent margins. Tectonophysics 219, 163–175. Wetherill G. W., Aldrich L. T., and Davis G. L. (1955) A40/ K40 ratios of feldspars and micas from the same rock. Geochim. Cosmochim. Acta 8, 171–172. Wilde S. A., Zhang X., and Wu F. (2000) Extension of a newly identified 500 Ma metamorphic terrain in North East China: further U–Pb SHRIMP dating of the Mashan Complex, Heilongjiang Province, China. Tectonophysics 328, 115–130. Wilde S. A., Valley J. W., Peck W. H., and Graham C. M. (2001) Evidence from detrital zircons for the existence of continental crust and oceans on the Earth 4.4 Gyr ago. Nature 409, 175–178. Williams M. L. and Jercinovic M. J. (2002) Microprobe monazite geochronology: putting absolute time into microstructural analysis. J. Struct. Geol. 24, 1013–1028. Williams M. L., Jercinovic M. J., and Terry M. P. (1999) Age mapping and dating of monazite on the electron microprobe: deconvoluting multistage tectonic histories. Geology 27, 1023–1026. Windley B. F., Allen M. B., Zhang C., Zhao Z. Y., and Wang G. R. (1990) Paleozoic accretion and Cenozoic redeformation of the Chinese Tien Shan Range, central Asia. Geology 18, 128–131. Windley B. F., Kro¨ner A., Guo J. H., Qu G., Li Y., and Zhang C. (2002) Neoproterozoic to Palaeozoic geology of the Altai orogen, Chinese Central Asia: new zircon age data and tectonic evolution. J. Geol. 110, 719–739. Wolf R. A., Farley K. A., and Silver L. T. (1996) Helium diffusion and low-temperature thermochronometry of apatite. Geochim. Cosmochim. Acta 60, 4231–4240. Wu F. Y., Jahn B. M., Wilde S., and Sun D. Y. (2000) Phanerozoic crustal growth: U–Pb and Sr–Nd isotopic evidence from the granites in northeastern China. Tectonophysics 328, 89–113. Yarmolyuk V. V., Kovalenko V. I., Kovach V. P., Budnikov S. V., Kozakov I. K., Kotov A. B., and Salnikova E. B. (1999) Nd-isotopic systematics of western Transbaikalian crustal protoliths: implications for Riphean crust formation in central Asia. Geotectonics 33, 271–286. York D., Hall C. M., Yanase Y., Hanes J. A., and Kenyon W. J. (1981) 40Ar/39Ar dating of terrestrial minerals with a continuous laser. Geophy. Res. Lett. 8, 1136–1138. Zartman R. E. and Doe B. R. (1981) Plumbotectonics—the model. Tectonophysics 75, 135–162. Zeitler P. K. (1988) Argon diffusion in partially outgassed alkali feldspars: insights from 40Ar/39Ar analysis. Chem. Geol. 65, 167–181. Zeitler P. K., Herczig A. L., McDougall I., and Honda M. (1987) U–Th– He dating of apatite: a potential thermochronometer. Geochim. Cosmochim. Acta 51, 2865–2868.

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Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 321– 348

3.11 Granitic Perspectives on the Generation and Secular Evolution of the Continental Crust A. I. S. Kemp and C. J. Hawkesworth University of Bristol, UK 3.11.1 INTRODUCTION 3.11.1.1 Granites and the Continental Crust 3.11.1.1.1 Granite terminology 3.11.1.1.2 Granites and crustal growth 3.11.2 BULK CRUSTAL COMPOSITIONS AND THE MANTLE-DERIVED PROTOLITHS 3.11.2.1 Assessing the Arc– Plume Balance 3.11.2.2 Episodicity in Crust Generation, and Its Implications 3.11.2.3 Evidence for Secular Changes in the Composition of the Continents 3.11.3 GENERATION OF HIGH SILICA CONTINENTAL ROCKS: I. THE ARCHEAN 3.11.3.1 Archean TTG Associations 3.11.3.1.1 Geochemical constraints on TTG genesis 3.11.3.1.2 Secular evolution in TTG composition 3.11.3.1.3 A combined model for TTG formation 3.11.3.2 Late Archean Granitic Rocks: The Onset of Major Intracrustal Melting 3.11.3.2.1 Generation of the Archean granitic types 3.11.3.2.2 Tectonic implications 3.11.4 GENERATION OF HIGH SILICA CONTINENTAL ROCKS: II. PROTEROZOIC TO PHANEROZOIC GRANITIC ASSOCIATIONS 3.11.4.1 The Importance of Intracrustal Melting, and the Mantle Connection 3.11.4.2 Peraluminous Leucogranites—Granites Derived from Old Crustal Protoliths 3.11.4.2.1 Conditions of leucogranite formation 3.11.4.2.2 Low-strontium leucogranites 3.11.4.2.3 High-strontium leucogranites 3.11.4.2.4 Relevance for crustal evolution 3.11.4.3 Granites Formed from Juvenile Mantle-derived Materials 3.11.4.3.1 A-type granites 3.11.4.3.2 Petrogenetic models for A-type granites 3.11.4.3.3 A-type case studies 3.11.4.3.4 Relevance for crustal evolution 3.11.4.4 Granites with Juvenile Mantle and Crustal Sources: The Lachlan Case Study 3.11.4.4.1 Mixing or unmixing? 3.11.4.4.2 Two contrasting granite types 3.11.4.4.3 Geochemical and isotopic comparisons 3.11.4.4.4 Constraints on the origin of the LFB granites 3.11.4.4.5 A petrogenetic model 3.11.4.4.6 Relevance for crustal evolution 3.11.5 CRUSTAL DIFFERENTIATION THROUGH EROSION AND SEDIMENTATION 3.11.5.1 The Link Between Rb/Sr Ratios and Granitic Sources 3.11.5.2 Resolving the Effects of Igneous and Sedimentary Processes 3.11.5.3 Relevance for Crustal Differentiation

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350 351 351 352 353 353 356 358 360 360 360 368 368 369 369 370 370 371 373 373 377 378 381 382 382 385 385 386 387 387 388 392 394 395 397 398 398 399 400

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3.11.6 SYNTHESIS AND AVENUES FOR FUTURE STUDY ACKNOWLEDGMENTS REFERENCES

3.11.1 INTRODUCTION Every geologist is acquainted with the principle of “uniformitarianism,” which holds that presentday processes are the key to those that operated in the past. But the extent this applies to the processes driving the growth and differentiation of the Earth’s continental crust remains a matter of debate. Unlike its dense oceanic counterpart, which is recycled back into the mantle by subduction within 200 Ma (see Chapter 3.13), the continental crust comprises buoyant quartzofeldspathic materials and is difficult to destroy by subduction. The continental crust is, therefore, the principal record of how conditions on the Earth have changed, and how processes of crust generation have evolved through geological time. It preserves evidence of secular variation in crustal compositions, and thus the way in which the crust has formed throughout Earth’s history. Exploring the nature and origin of these variations is the focus of this chapter. Continental rocks are highly differentiated, and so the crust is enriched in incompatible components compared to the primeval chondritic composition (see Chapter 3.01). Of these, water is perhaps the most relevant, both for the origin and evolution of life, and also for many models of crust generation and differentiation. Similarly, the mass of continental crust is just 0.57% of the silicate Earth, and yet it contains ,35% of the potassium (using the crustal composition estimates in Table 1). Continental rocks comprise the buoyant shell that was once thought to float on a basaltic substratum, inferred from the wide distribution of chemically similar continental flood basalts (von Cotta, 1858). The links with the adjacent oceans were perhaps unclear, “the greatest mountains confront the widest oceans” (Dana, 1873). Yet, it has long been argued that the rock that has the most similar composition to the average continental crust, andesite, may be generated by fractional crystallization of basalt (Daly (1914) and Bowen (1928); but see the contrary arguments of Kelemen (1995) and Chapter 3.18). The average age of the continental crust is old, almost 2 Ga, the processes of crust generation may have changed with time, and the early crust may have been generated and destroyed more rapidly than in more recent times (Armstrong, 1991; Bowring and Housh, 1995). The present consensus is that the modern Earth’s continental crust has a bulk andesitic composition (,61% SiO2), but it is lithologically and chemically stratified, such that a mafic

401 404 404

Table 1 Compositions of the continental crust and selected trace element ratios referred to throughout this chapter (all values are from Chapter 3.01). Major element oxides in wt.% and trace elements in ppm.

SiO2 TiO2 Al2O3 FeOt MnO MgO CaO Na2O K2O P2O5 ASI Mg# Ba Rb Sr Zr Nb Y Sc V Cr Co Ni Cu Zn Ga La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U Rb/Sr Rb/Ba Sr/Nd Eu/Eup Eu/Sr (Gd/Yb)N

Upper crust

Middle crust

Lower crust

66.60 0.64 15.40 5.04 0.10 2.48 3.59 3.27 2.80 0.15 1.03 46.7

63.50 0.69 15.00 6.02 0.10 3.59 5.25 3.39 2.30 0.15 0.85 51.5

53.40 0.82 16.90 8.57 0.10 7.24 9.59 2.65 0.61 0.10 0.75 60.1

60.60 0.72 15.90 6.70 0.10 4.66 6.40 3.07 1.81 0.13 0.85 55.3

628 82 320 193 12 21 14 97 92 17.3 47 28 67 17.5 31 63 7.1 27 4.7 1.0 4 0.7 3.9 0.83 2.3 0.3 2 0.31 5.3 0.9 17 10.5 2.7

532 65 282 149 10 20 19 107 76 22 33.5 26 69.5 17.5 24 53 5.8 25 4.6 1.4 4 0.7 3.8 0.82 2.3 0.32 2.2 0.4 4.4 0.6 15.2 6.5 1.3

259 11 348 68 5 16 31 196 215 38 88 26 78 13 8 20 2.4 11 2.8 1.1 3.1 0.5 3.1 0.7 1.9 0.24 1.5 0.25 1.9 0.6 4 1.2 0.2

456 49 320 132 8 19 21.9 138 135 26.6 59 27 72 16 20 43 4.9 20 3.9 1.1 3.7 0.6 3.6 0.77 2.1 0.28 1.9 0.3 3.7 0.7 11 5.6 1.3

0.25 0.13 11.85 0.72 0.0031 1.65

0.23 0.12 11.28 0.96 0.0050 1.50

0.03 0.04 31.63 1.14 0.0032 1.70

Bulk crust

0.15 0.11 16.00 0.93 0.0034 1.61

Introduction lower crust depleted in granitic components underlies an evolved middle and upper crust (Table 1; see Chapter 3.01 for a review of the compositional structure of the continental crust). But how has this been achieved? In the simplest terms, the continental crust grows magmatically, directly from the underlying mantle, and by the tectonic accretion of thicker portions of oceanic crust, such as island arcs (see Chapter 3.18) or ocean plateaus (see Chapter 3.16). However, these materials are predominantly basaltic, whereas the continental crust is andesitic, and neither most andesites, nor especially the large complement of silicic igneous rocks, can be extracted directly from a peridotitic source. Crust formation must therefore occur in at least two stages, first, melting of the mantle to produce basaltic magma, and second, either fractional crystallization or remelting of the basalt ultimately to produce the more evolved rocks of which the continental crust is dominantly composed. The products of the second stage may subsequently undergo further differentiation by one or more cycles of remelting (termed anatexis or “intracrustal melting”), which is integral in the stabilization of new continental crust, or by weathering and erosion at the Earth’s surface (see Kelemen (1995), and Rudnick (1995) and references therein). Uncertainties revolve around the tectonic settings under which these stages occur, the extent to which they may have been nonlinear or episodic through Earth’s history (Figure 1), and especially how the processes associated with each have changed with time. It is reasonably well established that the Earth has cooled considerably from the earliest Archean to the present day (Brown, 1986), and so it follows that the sites, rates and processes of magma generation, and consequently of crust formation, have changed too. Intracrustal melting is also likely to have been strongly influenced by changing thermal and geodynamic conditions in the Earth, governing the type of materials likely to undergo fusion, the compositions of derivative melts and the efficiency with which these segregate from their residues. One possible manifestation is the predominance of tonalitic– trondhjemitic batholiths in the Archean, in contrast to the potassium-rich granitic plutons of Phanerozoic orogenic belts (Martin, 1986; Drummond and Defant, 1990). Whether conventional plate tectonic scenarios operated in the Archean, and when these came into existence remains unclear, and this impacts on models of crust generation and recycling of crustal materials back into the mantle. Equally, isotopic studies have shown that the mantle itself has been progressively modified by the parent – daughter fractionation of certain radiogenic

351

100

n 16 % Continental crust

50

C

A

12

B

8

4

0

0 0

1

2 Ages (Ga)

3

4

Figure 1 The distribution of U/Pb ages from juvenile crust, with only one age (n) used for each supracrustal succession or pluton to avoid undue weighting from particularly well-studied areas (from Condie, 1998), combined with selected curves representing different models for the changing volumes of stable continental crust with time (A: minimum curve based on Nd isotopes in shales (authors’ compilation); B: from Collerson and Kamber (1999), and similar to Kramers and Tolstikhin (1997); and C: Taylor and McLennan (1985)).

isotopes (e.g., Sm –Nd, Rb – Sr, Lu – Hf) that occurs with the removal of crustal materials (e.g., O’Nions et al., 1979; Allegre et al., 1983; see also Chapter 2.13). Hence, the isotope composition of the magmatic flux across the Moho, and therefore of the potential building blocks for the continents, has also changed significantly during planetary evolution. To investigate these issues, and in keeping with the theme of secular change, the formative stages of the continental crust are addressed sequentially, from generation of the basaltic protolith, through to the subsequent intracrustal and supracrustal recycling processes leading to its differentiation. In this way, the subject material of the chapter will itself become progressively more differentiated, mirroring the secular compositional evolution of the continental crust. 3.11.1.1 Granites and the Continental Crust 3.11.1.1.1

Granite terminology

Given the felsic nature of the exposed continental crust, much of the debate has focused on the origins of relatively evolved rock types, and in particular granite sensu lato. This is appropriate, since regional mapping and seismic profiling demonstrates that granitic rocks, their volcanic equivalents, and metamorphic and sedimentary rocks of granitic composition, are easily the most abundant constituents of the upper continental

352

Granitic Perspectives

crust (Wedepohl, 1991). The rare earth element (REE) patterns of most post-Archean granitic rocks show the same distinctive features as that of the contemporary upper continental crust, notably enrichment in “light” REE (LREE, lanthanum to samarium), relative depletion in europium and flat “heavy” REE (HREE, holmium to lutetium) trends. Clearly, granitic plutonism may have a fundamental role in shaping the compositional structure of the crust, involving the ascent of buoyant, incompatible element-rich silicic magmas that leave behind residues that either reside in the deep crust or are recycled into the upper mantle. In addition, granitic batholiths formed by intracrustal melting potentially sample large areas of the middle to lower crust and may therefore provide compositional information on this inaccessible part of the Earth. For all of these reasons, granitic rocks, and their role in crustal evolution, are the emphasis of this chapter. However, the petrogenesis of granitic magmas is controversial, and thus many aspects of this role remain unclear. The literature on granites is complex, and daunting to nonspecialist and specialist alike. This is testament to the lithological and geochemical diversity of granites, which in turn reflects a large variation in their potential mantle and crustal source rocks, conditions of magma formation, evolutionary processes and their emplacement at differing crustal levels in a variety of tectonic environments. It follows that granite classifications should avoid genetic implications, and it is expedient to clarify some granite-related terms at the outset. For convenience, “granite” or “granitic rock” is here loosely defined to include plutonic rocks with more than 20% modal proportion of quartz relative to the feldspars, which encompasses the alkali feldspar granite through to tonalite lithological fields defined by the IUGS (Le Maitre, 1989). A widely used classification scheme is that of Chappell and White (1974), which subdivides granites into either I-type, sourced from an igneous or infracrustal precursor, and S-type, thought to have been derived by anatexis of sedimentary or supracrustal protoliths that have experienced a weathering cycle. There are also the volumetrically minor “A”-(alkaline, anhydrous, Table 2

and/or anorogenic) and “M” (mantle-like)-type granites, which are essentially subcategories of the I-types. The I– S scheme is widely used, but it is genetic, and it implies that granitic rocks have a simple crustal source, which can be consistently deduced from mineralogy and geochemistry. In practice, most granitic magmas contain contributions from multiple sources, and the extent to which the compositional features of high level granitic plutons faithfully reflect those of the protolith continues to be vigorously debated. Thus, we favor the use of mineralogical qualifiers (e.g., hornblende granite, cordierite granite) and simple chemical classification schemes, such as the alumina saturation index (ASI) of Shand (1927) that convey no implication for putative protoliths (Table 2). In the latter, granites are either peraluminous, metaluminous, or peralkaline. Peraluminous rocks have an excess of Al2O3 over that required to crystallize feldspar and have an alumina saturating phase, metaluminous rocks are just saturated in Al2O3, whereas peralkaline granites are deficient in Al2O3 relative to the alkalis.

3.11.1.1.2

Granites and crustal growth

Since granitic magmas cannot be extracted directly from the upper mantle, there can also be confusion over the language used in the context of models for the generation of new crust. Clearly granitic plutons represent new crust, at least in part, if they contain a component of juvenile mantle-derived material. In practice, this could occur in three ways: (i) fractional crystallization of primary basaltic liquid, (ii) mixing between partial melts of pre-existing crust and mantlederived magmas, or their differentiates, and (iii) by partial melting of young, mantle-derived mafic protoliths in the crust. Formally, the last case is by remelting within the crust, but it tends to be described in terms of crust generation, particularly when the mafic protolith separated from the mantle during the same tectono-thermal episode that subsequently generated the granite. The implication is that at the time of melting, the mafic protolith had neodymium, strontium, and/or lead isotope compositions indistinguishable from

Classification of granitic rocks according to the ASI of Shand (1927), where ASI ¼ molar [A12O3/(CaO þ Na2O þ K2O)].

Category

Chemical definition

Peraluminous Metaluminous

ASI . 1.0 ASI , 1

Peralkaline

ASI , ,1 Al/(Na þ K) , 1

Diagnostic mineralogy Aluminous biotite ^ muscovite, cordierite, garnet, Al-silicate Subaluminous biotite ^ hornblende, clinopyroxene accessory allanite, titanite, epidote Sodic pyroxene (aegirine) or amphibole (riebeckite –arfvedsonite)

Bulk Crustal Compositions and the Mantle-derived Protoliths those of its contemporary mantle source, given its short crustal residence time. An example would be the remelting of a lower crustal basaltic sill shortly after its emplacement by the underplating of a successive batch of basaltic magma. However, if the same basaltic sill was to be remelted, say 200 Ma later, then the silicic magma would be derived from older, pre-existing crust, and hence be regarded as crustal remelting. This distinction based on the age of the source is arbitrary, but it is useful to clarify how these terms are used.

3.11.2 BULK CRUSTAL COMPOSITIONS AND THE MANTLE-DERIVED PROTOLITHS The chemistry of the present-day continental crust is the frame of reference anchoring any discussion of secular evolution of crustal composition, or the processes of its formation. In this chapter we build on the considerable efforts of a number of scientists to develop an integrated geological and geophysical model for the continental crust (see Chapter 3.01 for a detailed treatment of the methodologies employed to infer crustal compositions). The preferred chemical estimates of the continental crust used throughout this chapter are listed in Table 1. The major element composition of the upper crust is well constrained, since this is the most accessible to sampling, both directly and via erosion and sedimentation, and different studies utilizing diverse databases have yielded remarkably similar results. SiO2 is ,61%, and Mg number (Mg#, molar Mg/(Mg þ Fe)) is ,55 for the bulk continental crust, and so it is more differentiated than any magma in equilibrium with the upper mantle. Trace-element abundances are more variable, as are estimates for the composition and proportion of the middle and lower crust. As we will see below, the latter are critical to any discussion of the mechanisms of crust formation and differentiation.

3.11.2.1 Assessing the Arc–Plume Balance Despite the elevated silica and potassium values compared to the bulk silicate Earth, the Mg# of the bulk continental crust is relatively high. Thus, the continents cannot be derived by known tholeiite fractionation trends, which involve marked iron enrichment, and the debate has focused on the extent that the continental crust can be generated by calc alkaline differentiation as seen in many orogenic andesite suites, or whether the parental magmas should be highmagnesium andesites (e.g., Kelemen, 1995). The latter may be generated by partial melting of

353

subducted oceanic crust, and subsequent interaction of those melts with the overriding peridotite in the mantle wedge (Kelemen, 1995; see also Chapter 3.18). Such rocks occur, e.g., in the western islands of the Aleutian arc, where the plate convergence rate is extremely low, and the wider context of this setting is difficult to evaluate. The significance of high-magnesium andesites for the evolution of the continental crust is explored by Kelemen et al. (2003). The distinctive trace-element features of the continental crust are it’s high incompatible element contents, including uranium, thorium, and potassium, and hence elevated heat production (see Chapter 3.02), the negative mantle normalized anomalies for niobium and tantalum, and high lead contents (Figure 2). Thus, it is characterized by low Nb/La, Ta/La, and Ce/Pb ratios relative to the oceanic crust and the upper mantle. Presumably, the initial stage in the development of continental crust involved the partial melting of primitive spinel and garnet peridotite to form mafic magmas in equilibrium with mantle olivine. Of primary mantle-derived magmas, low Nb/La and Ce/Pb ratios are only a feature of those related to subduction (Hofmann et al., 1986) (or where the mantle source has been overprinted by a subduction-derived component, Section 3.11.4.3.3) and so it is widely assumed that similar processes were responsible for the average composition of the continental crust (e.g., Arculus, 1999; Davidson and Arculus, 2001). However, there are also significant volumes of intraplate magmatism, and a number of authors have sought to assess the balance of intraplate and subduction-related magmatism in the generation of continental crust (e.g., Rudnick, 1995; Plank and Langmuir, 1998; Barth et al., 2000). Intraplate magmas tend to have smoother mantlenormalized trace-element patterns, without the anomalies that characterize subduction-related magmas (Figure 2(b)). Element ratios that reflect the presence or absence of such anomalies therefore provide the basis for assessing the contributions from magmas from the different settings (Figure 3). Other features of subduction settings are, first, that oceanic island arcs tend to have greater proportions of rocks with high MgO and low silica than their continental counterparts. Second, oceanic arc rocks tend to preserve more of the signature attributed to fluids released from the downgoing slab (e.g., high Ba/Th, U/Th, Sr/Nd), and a smaller subducted sediment contribution. This is most easily seen in the observation that high Ba/Th and U/Th ratios are best developed in the more depleted arc rocks that have relatively low thorium abundances (e.g., McDermott and Hawkesworth, 1991). In contrast, high Th/Ce ratios, which appear to be reasonably diagnostic of a sediment contribution,

354

Granitic Perspectives Continental crust

100 Upper crust Bulk crust Lower crust Middle crust 10

Mantle normalized Ba 1

Rb (a)

U Th

Nb K

La Ta

Pb Ce

P Sr

Zr Nd

Sm Hf

Ti Eu

Tb Gd

Ho Dy

Er Y

Yb Tm

Lu

Oceanic basalts 100

Bulk crust OIB

Sth. Sandwich Tonga-Kerm.

10

1

N-MORB Mantle normalized Ba

Rb (b)

U Th

Nb K

La Ta

Pb Ce

P Sr

Zr Nd

Sm Hf

Ti Eu

Tb Gd

Ho Dy

Er Y

Yb Tm

Lu

Figure 2 Primitive-mantle normalized minor and trace-element diagrams for: (a) the upper, middle, bulk, and lower continental crust (values from Table 1), and (b) oceanic and island arc basalts and the bulk continental crust (all normalizing values are from McDonough and Sun, 1995). The oceanic basalts (N-MORB, normal mid-ocean ridge basalt and OIB, ocean island basalt) are from Sun and McDonough (1989), whereas the arc basalts are from Turner et al. (1997) (Tonga-Kermadec arc) and Pearce et al. (1995) (South Sandwich arc).

tend to be better developed in relatively enriched arc rocks with higher incompatible element contents (Hawkesworth et al., 1997). Negative Nb – Ta anomalies are most prominent in the depleted island arc tholeiites, and these are not associated with relatively low titanium values (Figure 2(b); Pearce and Peate (1995) and see discussion of Figure 7(b)), as they are in most other magmatic suites (e.g., see Figure 5). The inference is that negative Nb – Ta anomalies may result from different processes in different magmatic suites (Section 3.11.3.1.1). The more evolved major element compositions in continental arcs are typically attributed to greater degrees of

differentiation within the thicker continental crust, often accentuated by crustal contamination (e.g., Davidson et al., 1991). Trace elements may be modified by contamination, but higher abundances are also often linked to larger contributions from subducted sediment, as proposed for lavas from the Aeolian arc, Italy (Ellam et al., 1989; Hawkesworth et al., 1993) or Lesser Antilles (e.g., MacDonald et al., 2000), not least because sediment fluxes at subduction zones are higher near the continents. In some cases this may be accompanied by smaller degrees of melting in the mantle wedge beneath thicker continental lithosphere.

Bulk Crustal Compositions and the Mantle-derived Protoliths 1.4 OIB

1.2

Nb/La

1

oceanic arcs average TTG

Primitive mantle

50%

average adakite average cont. arcs

MORB

0.8

20% LC

0.6

10% BC Archean

MC

0.4

UC

Adakites

Andes

0.2 0

0 (a)

10

20

30 Sr/Nd

40

50

60

0.3 oceanic arcs

UC

0.25

average TTG MC

average adakite average cont. arcs

Rb/Sr

0.2 BC

0.15

Andes Archean

0.1

0

Adakites

OIB Mantle MORB

0.05

0 (b)

10

LC 50%

20

30 Sr/Nd

10%

40

50

60

Figure 3 Mixing models for the formation of the bulk continental crust (BC), involving selected mantle-derived protoliths and based on key traceelement ratios (UC, upper crust; MC, middle crust; LC, lower crust). The crustal values for this, and subsequent figures are taken from Chapter 3.01, listed in Table 1 (a) Nb/La versus Sr/Nd and (b) Rb/Sr versus Sr/Nd. The end-members are the average OIB of Sun and McDonough (1989) and a primitive island arc basalt from Vanuatu (from Davidson and Arculus, 2001); other arc basalt data is from Davidson and Arculus (2001) and Turner et al. (1997). Crosses on the mixing curves correspond to 10% additions of the OIB component. Also shown for comparison are the primitive mantle and average MORB of Sun and McDonough (1989), the average andesite – dacite – rhyolite suite from the Andes and continental arcs (Drummond et al., 1996) and average Archean TTG compositions (from Martin, 1995 and Smithies, 2000). The adakite field is plotted from the sources quoted by Smithies (2000), and the average adakite composition is taken from Drummond et al. (1996). The bulk Archean crust is from Taylor and McLennan (1995).

In the context of this discussion, any assessment of the contributions of subduction and within plate magmatism is sensitive to the choice of the arc end-member. Many authors have selected

355

an oceanic arc as the subduction-related endmember, because these rocks contain less recycled material and hence reflect relatively more new material being added to the crust. Figure 3 illustrates variations in Nb/La, Sr/Nd, and Rb/Sr values for average crustal compositions, the primitive mantle, oceanic basalts (normal midocean ridge basalt (N-MORB); ocean island basalt (OIB)), and island arc lavas (island arc tholeiite (IAT)). Mixing results in near straight arrays on the plot of Nb/La –Sr/Nd, and it is striking that the average crustal compositions do not plot on any simple mixing curve between the island arc tholeiite, OIB or primitive mantle (see also Rudnick, 1995). If continental crust primarily consists of some mixture of magmas generated in island arc and intraplate (OIB-like), its bulk composition has since been modified by an additional process or processes. The additional processes are most likely to involve magmatic differentiation and weathering and erosion. However, Nb/La ratios are relatively robust in andesitic magmas, in the sense that they are not readily changed by fractional crystallization, crustal contamination, or weathering. In contrast, Sr/Nd ratios are reduced, and Rb/Sr ratios are increased towards the average crustal values by fractionation of plagioclase in igneous systems or during weathering and sedimentary transport, as strontium is soluble in seawater (we will return to this last point in Section 3.11.5.2). With respect to igneous processes, plagioclase may fractionate during magmatic differentiation, or it may be residual during partial melting, but in either case that fractionation must have taken place within the crust, in the stability field of plagioclase. These are longstanding conclusions, with the implication that strontium has been selectively recycled into the mantle, either as a component of magmatic cumulates (and/or residues) or altered oceanic crust (see Chapter 3.15) where strontium is concentrated within carbonate precipitates (Rudnick, 1995). Alternative solutions to this problem have been proposed, though all tend to founder on the difficulty of generating elevated Rb/Sr ratios directly from the upper mantle. For example, Ellam and Hawkesworth (1988) speculated that most of the continental crust was generated early in Earth’s history by processes that were different to those operating above modern active subduction zones, which inherently produce basalts with low average Rb/Sr ratios. In this scenario, the high Rb/Sr component is represented by primitive Archean tonalite – trondhjemite – granodiorite (TTG) associations and andesites, such as those of the 2.9– 2.5 Ga Zimbabwe craton, which have Rb/Sr ratios similar to bulk crustal values (Ellam and Hawkesworth, 1988). Unlike arc-related basalts, the TTG were generated from basaltic,

356

Granitic Perspectives

rather than peridotitic sources, at depths below the stability field of plagioclase (Section 3.11.3.1.1). Given the ancient average age of the continental crust (,1.8 Ga), there is little doubt that Archean magmatism contributed greatly to the formation of the continents. However, the average TTG composition, estimated from Archean provinces worldwide, has lower Rb/Sr ratios, and high Sr/Nd, than the bulk continental crust (Figure 3(b)). More recent discussions have explored whether the subductionrelated component might originate as partial melts of the subducted oceanic crust (i.e., “adakites,” Section 3.11.3.1.1), as seen in some recent arc settings like the western Aleutians (Kelemen et al., 2003; see also Chapter 3.18). Yet the Rb/Sr ratios of these lavas only approach those of the average crust in the most differentiated compositions (Kelemen, 2002, personal communication). In fact, the Rb/Sr ratios of global average adakite compositions do not significantly exceed those of arc tholeiites (Figure 3(b)). Thus, this model also requires removal of strontium-rich material from the continents. Additional evidence that the bulk composition of the continental crust has been modified is that it has a small negative europium anomaly (i.e., europium is depleted relative to neighboring elements samarium and gadolinium in chondrite normalized REE patterns; Table 1) since this is not a normal property of primary mantle-sourced magmas in any tectonic setting (Sun and McDonough, 1989; see Figure 3(b)). As with strontium, and in contrast to the other REE, europium is accommodated by plagioclase, and thus is sensitive to intracrustal differentiation processes. However, as will be demonstrated in Section 3.11.5.2, fractionation of europium from the other REE cannot be achieved by dissolution of plagioclase during modern weathering. With the caveat that crustal strontium and europium concentrations have not been underestimated, this suggests that the elevated Rb/Sr and Eu depletion of the bulk crust was accomplished by the removal of what were originally plagioclasebearing assemblages from the base of the crust. This can only occur in tectonically thickened crustal columns or overthickened arc edifices (.40 km thick), where the intrinsically buoyant plagioclase-rich material transforms with depth into denser garnet-rich granulite or eclogite, which is gravitationally unstable and can founder into mantle peridotite (Jull and Kelemen, 2001). This mechanism has been recently invoked to explain the paucity of early formed primitive (magnesium-rich) mafic to ultramafic cumulates in the Izu-Bonin (Arculus, 1999) and Aleutian (Kelemen et al., 2003) arc systems. In areas of continental collision, the same result

may be achieved by convective removal of lower crustal cumulates/residues, if the lithospheric mantle coupled to the lower crust has also been delaminated. Another point from Figure 3 is that crustal contamination processes will shift mixed OIB– IAT compositions towards the average crust, and it is noticeable that the Andean andesite plots close to the average crustal composition. However, crustal contamination processes do not provide information on the generation of new crust. Rather, the displacement of the crustal compositions away from the OIB– IAT array highlights that the residence time of highly incompatible elements at high levels in the crust is long compared with those of more compatible elements deep in the crust. Finally, given the envisaged decrease in mantle potential temperature from the Archean to the present day (Campbell and Griffiths, 1992; see also Galer and Mezger, 1998), it seems plausible that the relative supra-Moho fluxes from arc- and intraplate magmas may also have changed through time. Condie (1994, 2000) proposes that the proportion of plume-related oceanic plateau basalts in greenstone successions was greater in the Archean than in younger terranes, where arcrelated magmatism became dominant. One might therefore expect to see differences between the bulk composition of the Archean and postArchean crust with respect to key chemical indices or elemental ratios. However, we note in Figure 3 that the estimated bulk composition of Archean continental crust has slightly lower Nb/ La ratios than that of the post-Archean crust. Since high Nb/La ratios are taken to be indicative of an intraplate or plume component (and both are too often taken to be synonymous) in recent volcanic rocks (Hofmann et al., 1986; Condie, 1999), on face value this implies that the Archean crust had about the same, rather than the predicted higher, contribution from plume-derived magmas than that generated subsequently. An alternative, and preferred, interpretation is that the low Nb/La ratios reflect different processes in Archean rocks, and this is discussed further below (Section 3.11.3.1.1).

3.11.2.2 Episodicity in Crust Generation, and Its Implications It is increasingly recognized (e.g., Boher et al., 1992; McCulloch and Bennett, 1994; Stein and Hofmann, 1994) that the generation of the Earth’s continental crust may follow an episodic, rather than continuous, pattern through time, and this in turn bears upon unraveling the processes responsible for the extraction of crustal materials from the mantle. For example,

Bulk Crustal Compositions and the Mantle-derived Protoliths a striking feature of age – frequency diagrams for the continental crust, based on U – Pb zircon dates, are the peaks of major episodes of igneous activity, separated by periods of apparent quiescence (Condie, 1998, 2000) (Figure 1). Assuming that this is not simply an artifact of selective preservation of certain crustal segments, the periodicity lies at the heart of current debates on the origins of the Archean and Early Proterozoic continental crust, since it is taken as prima facie evidence by some that crust formation was dominated by major thermal pulses associated with the emplacement of mantle plumes. In these scenarios it is considered less likely that a global episodicity in crust forming events would result from the more continuous subduction-related processes (Albare`de (1998); but this is discussed further below). In contrast, no such peaks are observed in the last billion years, when it may be reasonably inferred that subduction-related magmatism was dominant. Gurnis and Davies (1986) suggest that this last point might reflect the enhanced erosion of elevated young crust, and thus preferential recycling of this material into the mantle. However, it seems unlikely that the crustal formation ages should be completely erased from the detrital zircon record. Precise U –Pb zircon analyses have further highlighted that major contemporaneous episodes spread across different continents, such as between 2.74 Ga and 2.66 Ga, and prompted speculation that these are the products of catastrophic turnover events in the Earth’s mantle (Stein and Hofmann, 1994; Condie, 1998). Such interpretations are consistent with models of mantle processes in a hotter Earth (Hill, 1993; Davies, 1995). However, Archean rocks are preserved in just 7.5% of the Earth’s surface (Goodwin, 1991), and so there is some uncertainty over the extent these age distributions and inferred episodicity are representative of major events in the Archean. Besides the geochemical constraints, discussed in Section 3.11.2.1, a challenge for models for episodic crustal growth based on mantle plumes is to explain the formation of the evolved silicic rocks that comprise the continental crust. This is because the overwhelming product of mantle plumes is basalt, either as juvenile oceanic crust such as oceanic plateaus (see Chapter 3.16), flood basalt provinces (see Chapter 3.03) or mafic underplates. How then does the episodicity of mantle plume activity translate into a similar periodicity in the generation of evolved continental crust? One possibility is that continental growth occurs by lateral amalgamation of thick, difficult-to-subduct oceanic plateaus against pre-existing continents (Albare`de, 1998). Here, the basaltic materials are ultimately

357

reprocessed into felsic continental crust by intracrustal differentiation, involving anatexis and chemical weathering (Albare`de, 1998), and typically this has to be in short time periods to ensure the mantle-like initial isotope ratios of many of the felsic rocks. Critically, age peaks in the Archean and the Early Proterozoic in Figure 1 (the “super-events” of Condie (1998)) are observed in both mantle-derived volcanic rocks and in higher silica plutons that are probably crustal melts. Such close temporal links are clearly an important constraint on models of crust generation, and for each pulse of igneous activity there appears to be a consistent delay of 50 –80 Myr between the eruption of mantle-derived volcanics and the intrusion of syn-tectonic granites. Recent examples of felsic magmatism associated with plume activity tend to be temporally associated with large volumes of mafic rocks. The A-type granites of intraplate settings are very different from Archean TTG suites, and their compositions reflect the alkaline to enriched nature of the associated basalts, and the distinctive volatile component (Section 3.11.4.3.2). As we emphasize in Section 3.11.3.1, a number of continental flood basalt provinces are markedly bimodal in their silica distributions, similar to many Archean terranes. However, the felsic rocks of flood basalt provinces tend not to have been generated in equilibrium with residual garnet, which characterizes the Archean TTG (Section 3.11.3.1.1). The key element may therefore be the estimated time lag of 50 – 80 Myr between the eruption of mantle-derived volcanics and the intrusion of syn-tectonic granites, since it highlights that the mafic and felsic components of Archean crust were generated under different conditions. The tectonic settings may therefore have been different, or at least a particular setting would have evolved in character over the inferred 50 –80 Myr. An alternative to oceanic plateau accretion is that continental growth reflects a transient surge of accelerated subduction, this being a byproduct of enhanced heat flow in the upper mantle attending plume activity. The intrusion of syn-tectonic granites is thought to form through collision or accretion of the juvenile crustal fragments (Condie, 1998). The emplacement of silicic rocks, and thus the growth of continental crust, would, in this model, reflect closure of oceanic basins by subduction. Subduction rates, and thus the volumes of arc magmas, are likely to increase dramatically during periods of high plume flux in the upper mantle, since this will increase the generation rate of juvenile oceanic crust along ridge systems (Condie, 2000). Furthermore, the subducted crust will be younger, which thermally favors slab melting and the production of the

358

Granitic Perspectives

felsic TTG that comprise large parts of Archean cratons and the earliest continents. This model has the attraction of coupling mantle plume activity and subduction in the formation of new continental crust. Another important question is why the first age peak only on Figure 1 occurred at ,2.7 Ga and not before, since ancient detrital zircon ages demonstrate the existence of continental crust dating back to ,4.4 Ga (Wilde et al., 2001). Several authors (e.g., Stein and Hofmann, 1994; Condie, 1998, 2000) propose that the mantle plumes initiating the continent-forming “superevents” are triggered by catastrophic “avalanches” of subducted eclogitic slabs that have accumulated at the 660 km seismic discontinuity. In this scenario, the absence of age peaks prior to ,2.7 Ga might reflect the efficient recycling of subducted oceanic crust into TTG magmas in the hotter early Archean mantle. Continental materials were also likely to be more rapidly returned to the convecting mantle during this period (Armstrong, 1991), and thus have a lower survival rate. Conversely, the 2.7 Ga episode may signal the onset of subduction and the emplacement of large crustal volumes during a short time interval. Clearly, the genesis of the TTG holds the key to understanding the formation of Earth’s first continents, and we shall address this problem in Section 3.11.3.

3.11.2.3 Evidence for Secular Changes in the Composition of the Continents There are clear differences in the common igneous rock types present in Archean and postArchean regions, and hence in the distribution of major and trace elements, as seen in the preponderance of bimodal silica distributions in Archean cratons compared with those in younger terranes generated along convergent continental margins. Allied with the pattern of episodicity apparent in Figure 1, and the progressive cooling of the Earth, this suggests that we might also expect to resolve some secular changes in crustal compositions. Potentially, the best evidence for this is from detrital continental sediments, since it is generally agreed that they provide good estimates of the average post-Archean upper crust, at least for elements, like scandium, thorium, yttrium, and the REE, that are not soluble in water (Taylor and McLennan (1985, 1995); see also the review by Rudnick and Gao (Chapter 3.01). However, we need to be aware that the extent to which any differences in sedimentary rock geochemistry through time manifest changes in the composition of the bulk continental crust is much more difficult to assess.

Taylor and McLennan (1985) established that compared to post-Archean shales, Archean clastic sedimentary assemblages have systematically lower silicon, potassium, thorium, and scandium contents and higher magnesium, calcium, sodium, chromium, and nickel. This has been largely corroborated by Condie (1993), using cratonic shales, though lower calcium, sodium, and strontium in Archean shales compared to their younger counterparts implies that chemical weathering was more intense in the CO2-rich atmosphere of the early Earth (Condie, 1993). The differences in compositions between Archean and younger sedimentary rocks documented by Taylor and McLennan (1985) grossly mimic the chemistry of the dominant igneous rocks emplaced during these periods, notably with the preponderance of sodic TTG and magnesium-rich komatiites in the Archean, compared to the more siliceous and potassium-rich granites in younger orogens (Sections 3.11.3 and 3.11.4 respectively). Even allowing for the removal of a large basaltkomatiite component by erosion, Archean shales also have lower titanium, niobium, and tantalum than their post-Archean counterparts (Condie, 1993), which seems to reflect derivation from a provenance dominated by TTG, which are depleted in these elements (see Figure 5(a)). As such, Nb/La ratios are lower in Archean shales (0.36) and the Archean upper crust (0.28), compared to post-Archean shales (0.44 – 0.50) and contemporary upper crust (0.39 – 0.46; all values from Condie (1993)), though such differences are less apparent in bulk crustal compositions (Figure 3). The implications of this for crustal growth processes were highlighted in the above section. However, interpretation of the REE data has provoked more controversy, largely reflecting the choice of Archean sediment from which the estimates derive. Taylor and McLennan (1985, 1995) suggest that, although considerable variability exists, low-grade graywackes of Archean greenstone belts exhibit less LREE enrichment, greater fractionation of the HREE (i.e., extend to higher Gd/Yb ratios) and no negative europium anomaly compared to post-Archean shales, which have marked negative europium anomalies. These features are in turn reflected in contrasting chondrite-normalized REE patterns for the Archean and post-Archean crust (Figure 4; and see Taylor and McLennan (1985), figure 2)). As we shall see in Section 3.11.3.1.1, steep REE patterns lacking significant europium anomalies characterize the Archean TTG. Europium is preferentially partitioned into plagioclase feldspar, and so Taylor and McLennan (1985, 1995), conclude that residual plagioclase was much more of a feature of intracrustal differentiation processes in

Bulk Crustal Compositions and the Mantle-derived Protoliths Upper continental crust 100 Archean UC (Condie 1993) Archean UC (T&M 1995) modern UC (R&G)

10

Chondrite normalized La Ce Pr Nd

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Figure 4 Comparison between the Archean upper crust (UC) and the contemporary upper crust of Rudnick and Gao (see Chapter 3.01) (R&G) with respect to chondrite-normalized REE patterns (normalizing values from Anders and Grevesse, 1989). Archean data are from Condie (1993) and Taylor and McLennan (1995) (T&M 1995).

the post-Archean, compared with the Archean crust. The implication is that intracrustal melting at depths shallow enough to be in the plagioclase stability field was less common in the Archean. Since strontium is also preferentially partitioned into plagioclase relative to rubidium, it may also be inferred that the post-Archean upper crust will have a higher Rb/Sr ratio, and this is evident in Figure 3. The MREE to HREE are fractionated by residual garnet, and overall the Archean and post-Archean upper crusts have similar MREE to HREE patterns. However, the Archean patterns are attributed to variable mixing between flat MREE to HREE greenstone belt volcanic rocks with highly fractionated patterns from the TTG, typically in a ratio of two parts basalt to one part TTG (Taylor and McLennan, 1985). In contrast, the similar patterns in the post-Archean upper crust are taken as evidence of the lack of residual garnet in younger crust differentiation processes, since rocks like the Archean TTG are rare. Alternatively, Gibbs et al. (1986) and Condie (1993) suggest that rather than reflecting the absence of intracrustal differentiation prior to 2.7 Ga, the secular variation in the REE is simply an artifact of comparing Archean graywackes with post-Archean shales that have different tectonic settings. Refuting this, Taylor and McLennan (1995) point out that the negative europium anomaly, and thus the signature of crustal differentiation, persists in modern graywackes of all tectonic settings and is present even in forearc turbidites and pelagic clays remote from the continents. It, therefore, seems inescapable that if mid- to shallow-level intracrustal differentiation were widespread in the Archean, the characteristic europium

359

depletion would be recorded by greenstone belt sedimentary sequences. Gibbs et al. (1986) and Condie (1993) further argue that cratonic (passive margin) shales are more representative of the Archean upper crust than active margin greenstone belt graywackes. Condie (1993) emphasizes that cratonic Archean shales posses negative europium anomalies that overlap statistically with those of post-Archean shales, diminishing the difference between the Archean and post-Archean upper crust with respect to the REE. Nevertheless, even if this approach is justified (see the contrary arguments of Taylor and McLennan (1995)), the Archean upper crust of Condie (1993) still has less relative depletion in europium than the post-Archean upper crust (Figure 4). It seems therefore that these differences are real, and relate to contrasting behavior of major phases (plagioclase, garnet, and possibly amphibole) during magmatism and crustal differentiation in the Archean and post-Archean periods. The lower crust is much more difficult to sample, and this in turn limits the accuracy of estimates of bulk continental crust in the Archean and younger terranes. The common approaches, summarized by Rudnick and Gao (Chapter 3.01), include studies of xenolith populations in alkali basalts and kimberlites, high-grade metamorphic terranes, and seismic velocity structure. In general, while the evidence for crustal melting in the generation of granites is overwhelming (Section 3.11.4), residual assemblages after partial melting are rarely sampled. Nonetheless, the best estimate of the lower crust is that it is mafic and comprises the residua left after the extraction of the granodioritic upper crust, perhaps subsequently underplated by basaltic magmas (Taylor and McLennan, 1985, 1995). We note that the current lower crustal composition plots near the arc basalt-OIB mixing lines on Figure 3, which is consistent with a mafic component generated in both subduction-related and intraplate settings. There is evidence from xenolith and seismic imaging studies that the Archean continental crust is thinner, and generally lacks the basaltic underplating and development of mafic granulitic lower crust observed in postArchean regions (see Chapter 3.01). One possible reason for this is that the mafic complement was part of the original crustal architecture but has since been lost to the convecting mantle, as suggested by Nui and James (2002) for the Kaapvaal Craton (southern Africa). Alternatively, mafic lower crust may never have existed in Archean cratons, reflecting a fundamentally different style of crustal generation during this period (Section 3.11.3.1.1).

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Granitic Perspectives

In summary, sediments and xenolith suites preserve a record of gross secular change in the compositional structure, and perhaps the bulk composition of the continental crust. Given that evolved igneous rocks are the building blocks of the continents, an important step towards understanding the mechanisms driving these changes is to examine silicic igneous activity in the Earth through time, and this is the focus of the following sections. Another aim is to elucidate the tectonic “engine” controlling crustal generation and evolution. It is also necessary to unravel the relative contribution of other differentiation processes, such as weathering and erosion, to the chemistry of the continental crust, and this is treated separately at the end of the chapter. By way of introduction, average analyses of silicic igneous rocks from various time periods are listed in Table 3, and normalized to the bulk continental crust in Figure 5. Key chemical features are highlighted in bivariate elementelement and ratio –ratio plots (Figure 6). Two aspects are immediately evident. First, there are clear differences between Archean tonalites and most Phanerozoic plutonic rocks, particularly with regard to Na/K and the REE (Figure 6). Second, none of the Phanerozoic granite types mimic the current upper crustal composition for the full range of trace elements considered, with the greatest discrepancies shown by the most primitive (oceanic arc) and felsic (leucogranites) rocks. This suggests that other igneous and/or sedimentary processes must be invoked to account for the generation of the uppermost granodioritic crust.

3.11.3 GENERATION OF HIGH SILICA CONTINENTAL ROCKS: I. THE ARCHEAN 3.11.3.1 Archean TTG Associations The implications from neodymium model ages on upper crustal rocks is that most crust generation occurred relatively early, such that at least 50% of the continental crustal mass was emplaced by the end of the Archean (Taylor and McLennan, 1985; Figure 1, and Chapter 3.10). Although little of this embryonic crust is now exposed at the Earth’s surface, bimodal silica distributions in Archean terranes show that the massive igneous activity during this period produced two contrasting suites, high-magnesium basalts (including basaltic komatiites and komatiites) and the relatively felsic TTG. With an average of ,70% SiO2 (Martin, 1995) the latter are among the oldest high silica igneous rocks preserved and comprise the earliest continental nuclei (e.g., the ca. 4.0 Ga Acasta

gneisses of the western Slave Province, Canada, Bowring and Williams, 1999). Although the tectonic setting for Archean magmatism remains unresolved, it is noteworthy that in the recent geologic past bimodal silica suites are a feature of extensional and plume-related settings, as in continental flood basalt-rhyolite provinces like the Parana-Etendeka Province (Garland et al., 1995; Peate, 1997) or the Karoo Igneous Province (Duncan et al., 1984), rather than of convergent plate margins.

3.11.3.1.1

Geochemical constraints on TTG genesis

The TTG are different from most of the evolved plutonic rocks of Proterozoic and Phanerozoic orogenic belts (Figures 5 and 6; Table 3). This is most evident in their typically high Na2O and Al2O3 and low K2O contents, and steep REE patterns, with strongly elevated strontium and LREE, and depleted HREE, yttrium, and scandium. In contrast to most younger plutons, the most “primitive” TTG also commonly lack negative europium anomalies, and although many TTG in Figure 6 do have small negative europium anomalies, these are generally smaller than that of the bulk continental crust. Strontium and neodymium isotopic ratios of the TTG are primitive and approach mantle values at the time of crystallization (Martin, 1995), though some TTG suites have apparently assimilated older crustal materials (e.g., those of the Superior Province, Canada; Whalen et al. (2002)). This striking combination of features is generally attributed to derivation from a hydrated (meta)basaltic source rock at depths below the stability field of plagioclase (i.e., .40 km) but in equilibrium with residual garnet, to account for the low HREE contents (Martin, 1986, Martin, 1995; Luais and Hawkesworth, 1994; Rapp, 1997; Wyllie et al., 1997). Isotopic considerations demand a short crustal residence time for the basaltic protolith prior to melting, to preserve the mantle-like 87Sr/86Sr and 143Nd/144Nd ratios, and thus emplacement of the TTG constitutes growth of new continental crust. However, the tectonic implications are unclear because partial melting may have taken place either in subducted slabs, similar to destructive plate margins at the present day (Drummond and Defant, 1990; Martin, 1995; Condie, 1998), or in underplated basalt beneath thickened crust or oceanic plateau (Kro¨ner, 1985; Smithies, 2000; Whalen et al., 2002). The underplated basalts may have been generated in response to mantle plume activity, or above old subduction zones.

Table 3 Representative analyses of the granite types referred to throughout this chapter and plotted in Figure 5 (see Figure 5 caption for data sources). Sample

SiO2 TiO2 Al2O3 Fe2O3 FeO FeOt MnO MgO CaO Na2O K2 O P2O5 LOI Sum ASI Mg# Ba Rb Sr Zr Nb Y Sc V Cr Co Ni Cu Zn Ga La Ce Nd

TTG (n ¼ 355)

Archean calc-alkaline Type 1

Type 2

69.79 0.34 15.56

70.00 0.40 14.63

71.88 0.23 14.69

2.81 0.05 1.18 3.19 4.88 1.76 0.13

2.72 0.05 0.84 2.28 3.89 3.58 0.17 0.78 98.39 1.015 35.5 1300 117 479 218 12 21 4.7 28 26 13 12 19 59 18 71 133 45

1.65 0.03 0.48 1.69 4.45 3.69 0.08 0.81 98.79 1.021 34.1 1210 125 455 142 9 7 1.8 15 73 14 12 13 41 17 42 71 29

99.56 0.989 42.8 690 55 454 152 6.4 7.5 4.7 35 29 14

32 56 21.4

Cont. arc. PRB (n ¼ 323)

Oceanic arc (n ¼ 8)

64.63 0.65 15.94 1.20 3.19 6.06 0.08 2.15 5.10 3.62 1.95 0.13

68.10 0.48 14.99 1.77 2.28 4.32 0.10 1.29 4.36 3.98 1.38 0.09

98.51 0.919 38.7 641 60 375 139 6.7 19 14 85 47

98.73 0.938 34.8 311 19 267 112 1.3 22 14 63 2 18 2 35 50 14.8 6.3 16.3 11.3

13 10 76 18.3 16 35 14

GRC Leucogran. (n ¼ 8) 74.59 0.05 14.63 0.14 0.65 1.24 0.02 0.29 1.74 3.03 3.97 0.105 0.89 99.98 1.176 29.7 1200 125.8 283.1 16.6 3.73 20.1 5 45 35 15 11 19 56 8.86 18.56 7.39

Himalayan Leucogran. (n ¼ 13)

A-type Padthaway (n ¼ 6)

74.16 0.09 14.74

75.93 0.17 11.91

0.90 0.02 0.17 0.70 3.91 4.43 0.17 0.69 98.22 1.179 25.2 223 229 124 35.7 9.8 10.8

1.93 0.07 0.12 0.40 3.63 4.91 0.02 0.29 97.14 0.992 9.7 76 132 7 330 22 49 3 2 5 3

10.8 23.2 10.8

20 147.5 226.93 118.41

A-type LFB

72.06 0.38 12.43 1.61 1.55 2.95 0.08 0.40 0.93 3.94 4.13 0.09 2.02 97.51 0.983 19.5 725 159 95 460 24.5 79 15.2 8 2 4 1 4 125 20.2 64 132 64

LFB Hbl granite (n ¼ 1074)

LFB Crd granite (n ¼ 704)

69.50 0.41 14.21 1.01 2.22 3.12 0.07 1.38 3.07 3.16 3.48 0.11

70.91 0.44 14.00 0.52 2.59 3.06 0.06 1.24 1.88 2.51 4.09 0.15

98.51 0.977 44.1 519 164 235 150 11 31 13 57 20 10 8 9 48 16 31 66 23.7

98.24 1.169 41.9 440 245 112 157 13 32 11 49 30 10 11 9 59 18 27 61 21.9 (continued )

Table 3 Sample

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U Rb/Sr Rb/Ba Sr/Nd Eu/Eup Eu/Sr (Gd/Yb)N

(continued). TTG (n ¼ 355)

3.3 0.92 2.2 0.31 1.16

Archean calc-alkaline Type 1

Type 2

7.8 1.56 5.1 0.73 3.2

4.4 0.79 3.4 0.41 2.2

Cont. arc. PRB (n ¼ 323)

Oceanic arc (n ¼ 8)

4.1 0.93 4.0

GRC Leucogran. (n ¼ 8) 2.05 1.11 2.34

Himalayan Leucogran. (n ¼ 13) 3 0.63

A-type Padthaway (n ¼ 6) 13.52 0.43 7.52

0.59

6.9 1.6 0.12 0.08 21.21 1.037 0.0020 3.30

1.55 0.24 5.8 0.93 23 21 2.4 0.24 0.09 10.64 0.751 0.0032 2.72

0.51 0.09 4.6 0.29 28 22 3.3 0.27 0.10 15.69 0.620 0.0017 5.51

14.9 2.42 12.4

LFB Hbl granite (n ¼ 1074)

LFB Crd granite (n ¼ 704)

4.9

5.8

19 20 5 0.70 0.31 9.91

27 19 5 2.19 0.56 5.11

0.56 3.02

6.48

1.80

3.43

1.2 0.55 0.12 4.5 0.71

A-type LFB

10 7.2 1.5 0.16 0.09 26.78

2.25 0.50 3.3 0.88 5 0.83 0.37 0.07 0.06 23.73 0.692 0.0035 1.47

3.2 1.67 0.24 2.29 0.45 55.8 3.42 2.09 0.44 0.10 38.31 1.537 0.0039 1.15

1.29 0.18 80 5 1.84 1.02 11.48 0.61 0.0051

3.36

17 22 2.9 19.55 1.73 0.06 0.130 0.0614 1.85

1.1 8.2 1.27 11.1 2.9 22 19.7 5 1.67 0.22 1.48 0.541 0.0254 1.25

TTG, Archean tonalite-trondhjemite-granodiorite; PRB, Peninsula Ranges Batholith; GRC, Glenelg River Complex; LFB, Lachlan Fold Belt; Hbl, hornblende granite; Crd, cordierite grantie.

Generation of High Silica Continental Rocks: The Archean

363

intracrustal recycling processes in TTG genesis is highlighted by Whalen et al. (2002), and references therein. One argument is that the TTG are chemically analogous to modern adakite lavas, which are confined to arc settings and largely attributed to melting of

At present there are a number of competing lines of evidence in support of the TTG being formed by remelting at the base of thickened crust, and in the subducted slab. Support for the subduction model is reviewed by Martin (1995), whereas evidence for the role of

10 Archean plutons

TTG granite CA1 granite CA2

1

Bulk crust normalized Ba 0.1

Rb

U Th

Nb K

La Ta

Pb Ce

Nd Sr

Zr P

Sm Hf

Gd Eu

Y Ti

Yb

(a) 10 Convergent margin granites oceanic arc PRB

1

Bulk crust normalized Ba 0.1

Rb

U Th

Nb K

La Ta

Pb Ce

Nd Sr

Zr P

Sm Hf

Gd Eu

Y Ti

Yb

(b)

Figure 5 Multi-element diagrams of the different silicic igneous rock types discussed in the text compared to estimates of the upper continental crust (unornamented line), with all compositions are normalized to the bulk crust (values in Table 1). (a) Average of 355 Archean TTG (from Martin, 1995) and average Late Archean “calc-alkaline” (CA) granites (from Sylvester, 1995), (b) convergent margin granites; continental arc, Peninsula Ranges Batholith (PRB) California, from Silver and Chappell (1988), oceanic arc (New Britain) from Whalen (1985), (c) leucogranites; Glenelg River Complex (GRC, southeastern Australia, Kemp, 2001), Himalayan leucogranite from Inger and Harris (1993), (d) A-type granites; average Padthaway Suite granite, southern Australia (from Turner et al., 1992a), Gabo Island metaluminous granite, Lachlan Fold Belt (LFB, southeastern Australia, from Collins et al. (1982), peralkaline granite, LFB (Wormald and Price, 1988), (e) average hornblende (n ¼ 1074) and cordierite (n ¼ 704) granites, LFB (SE Australia), from Chappell and White (1992). Note the higher rubidium, but much lower Sr content of the cordierite granites, consistent with a greater component of weathered (meta)sedimentary material in the source region (Chappell and White, 1992).

364

Granitic Perspectives Leucogranites 10

GRC Himalayan

1

0.1 Bulk crust normalized Ba

0.01

Rb

U Th

Nb K

La Ta

Pb

Nd

Ce

Sr

Zr

Sm

P

Gd

Hf

Y

Eu

Ti

Yb

(c) 1,000

A-type granites

100

10

1

0.1 Padthaway LFB metaluminous 0.01

Bulk crust normalized Ba

0.001

Rb (d)

U Th

LFB peralkaline

Nb K

La Ta

Pb

Nd

Ce

Sr

Zr P

Sm Hf

Gd Eu

Y Ti

Yb

10 Lachlan Fold Belt granites

1

hornblende granite cordierite granite

Bulk crust normalized Ba 0.1

Rb

U Th

Nb K

La Ta

Pb Ce

Nd Sr

Zr P

(e)

Figure 5

(continued ).

Sm Hf

Gd Eu

Y Ti

Yb

Generation of High Silica Continental Rocks: The Archean

365

10

7 Archean TTG A-types

6

Himalayan LG

Cordierite granites

GRC LG

Eu/Eu*

K2O (wt.%)

5 4

1

Bulk crust

3 2

Hornblende granites

1

0.1 0 0 (a)

1

2

3 4 5 Na2O (wt.%)

6

7

8

0.1 (b)

1 (Gd/Yb)N Archean TTG Hornblende granites Cordierite granites A-types Himalayan LG GRC LG

100

10

Rb/Sr

10

1 Bulk crust 0.1

0.01

0.001 0.1 (c)

1 Eu/Eu*

10

Figure 6 The reference igneous suites shown on plots of (a) K2O versus Na2O (in weight percent), (b) Eu/Eup versus (Gd/Yb)N, and (c) Rb/Sr versus Eu/Eup (LG, leucogranite; GRC, Glenelg River Complex). The subscript N refers to the abundance of the element in the rock normalized to that of the chondritic composition recommended by Anders and Grevesse (1989). Eu/Eup is a measure of the normalized p abundance of Eu relative to that of Sm and Gd (the “Eu anomaly”) and is computed according to Eu/Eup ¼ EuN/ [(SmN)·(GdN)]. Data sources are: TTG, from the references quoted by Smithies (2000), and including Kamber et al. (2002) and Whalen et al. (2002); LFB hornblende and cordierite granites are from Anderson (1998), Chappell et al. (1991), Chappell and White, 1992; Chen et al. (1989, 1990), Elburg (1996b,c), Griffin et al. (1978); Hine et al. (1978), Kemp (unpublished data), Maas et al. (1997), Munksgaard (1988), Soesoo (2000), White and Chappell (1988, 1989), White et al. (1977); A-type granites, Collins et al. (1982), King et al. (1997, 2001), Turner et al. (1992a), Whalen et al. (1987), Wormald and Price (1988), and some data from the Geoscience Australia Rockchem repository; Himalayan leucogranites, Ayres and Harris (1997), Crawford and Windley (1990), Inger and Harris (1993); France-Lanord and Le Fort (1988), Le Fort et al. (1987), Searle et al. (1997), Vidal et al. (1982); Glenelg River Complex leucogranites, Kemp (2001).

the downgoing oceanic lithosphere (see reviews by Defant and Drummond, 1990; Drummond and Defant, 1990; Peacock et al., 1994; Drummond et al., 1996; Martin, 1999). Yet as emphasized by Smithies (2000), sodic plutons that resemble the TTG, at least in terms of major elements, are also a minor component of overthickened continental arcs, such as in the Mesozoic batholiths of the western American Cordillera (Atherton and Petford, 1993; Tepper

et al., 1993; Petford and Atherton, 1996), New Zealand (Muir et al., 1998), and Antarctica (Wareham et al., 1997), where they are thought to derive by fusion of a juvenile underplate. A key difference between these two geodynamic models is that melts from subducted oceanic crust migrate through mantle en route to the crust, whereas those from underplated basalts do not. Many TTG, and adakites (e.g., Stern and Kilian, 1996), have higher MgO and Ni

366

Granitic Perspectives

contents than experimental melts of basalt (Martin, 1999; Martin and Moyen, 2002). These features have therefore been taken to reflect interaction of TTG magmas with peridotite in the mantle wedge (Drummond et al., 1996; Rapp et al., 1999; Martin and Moyen, 2002), and hence to be supportive of the subduction zone model. Like adakites, TTG extend to lower Nb/Ta and higher Zr/Sm ratios than the primitive mantle and oceanic basalts. Recent experimental work suggests that these characteristics reflect the presence of residual amphibole, and that those amphiboles had low Mg numbers of 40 – 50 (Foley et al., 2002). This requires melting of low magnesium amphibolite, rather than eclogite or even amphibole-bearing peridotite. Such fractionated basaltic rocks are presumed to be more typical of the upper parts of oceanic crust generated and hydrated at mid-ocean ridges, and later consumed by subduction, rather than the lower crust (Foley et al., 2002), where dry, magnesium-rich cumulates and eclogite are expected to predominate (Saunders et al., 1996).

Primitive TTG

0.04 DTh = 0 0.01 3 2 DTi = 1

0.06

Primitive mantle

0.05

OIB

0.1 4 TTG source rocks (F = 10%)

Ta/La

10

Th (ppm)

A striking feature of the TTG is that they have higher incompatible element contents and markedly lower TiO2, Y, and Yb contents that their inferred basaltic protoliths, and these can be used to constrain the melt generation models. Figure 7(a) summarizes the arguments for the incompatible elements thorium and TiO2 by illustrating the thorium and TiO2 contents of the inferred source rocks for different values of partition coefficients, assuming that an average TTG reflects 10% melting. Critically, to reproduce the high thorium contents of primitive TTG (with , 0.4% TiO2) simple modeling requires the basaltic protolith also to have moderately high thorium contents (.1 ppm) at low TiO2 for reasonable degrees of melting. At higher degrees of melting, these source thorium contents would also have to be higher, and basalts with 1 ppm thorium contents lie outside the range of contemporary depleted MORB. However, they plot within the upper envelope of the field for Archean basalts and komatiites, and are also similar to the Archean basalt used in the modeling of Foley et al. (2002). Many of these high thorium komatiites have Th/Ta ratios

1 E-MORB

0.04 0.03

Cont. crust Idaho Bath. 1.6

Komatiites and basalts

0.02

av. TTG 0.8

0 (a)

1

2 TiO2 (wt.%)

3

4

0

(b)

GLOSS

South sandwich arc

ZIMB. 1.0

0.01

N-MORB

0.1

Archean basalt

Aeolian Islands

10

Tonga arc

100 Ti/Zr

Figure 7 Batch melting models illustrating the possible generation of the Archean TTG from basaltic protoliths (cont. crust is the bulk continental crust from Table 1). (a) Plot of Th versus TiO2, showing the source composition (gray diamonds) modeled as a function of different bulk partition coefficients required to reproduce the Th and titanium contents of primitive TTG (12 ppm Th, 2,500 ppm Ti) by 10% batch melting. Having higher degrees of melting increases the implied Th content of the source. Various oceanic basalts (from Sun and McDonough, 1989; E-MORB, enriched mid-ocean ridge basalt) and the compositional field of Archean basalts and komatiites are shown for reference. (b) Comparison between oceanic arc basalts (,53% SiO2), adakites (closed diamonds), TTG, the primitive mantle, and modeled partial melts (gray diamonds) from an Archean basalt (closed circle) at different pressures (0.8 GPa, 1.0 GPa, and 1.6 GPa) in terms of Ta/La versus Ti/Zr (ZIMB ¼ least evolved TTG from Zimbabwe, Luais and Hawkesworth, 1994). The average TTG (n ¼ 355) is from Martin (1995). GLOSS represents the average global subducted sediment composition estimated by Plank and Langmuir (1998). Other data sources are as for Figures 2 and 3, with the Aeolian Island field from Ellam et al. (1989) and the Idaho Batholith field from Clarke (1990). The partial melts were modeled by 10% batch melting of the average Archean Sula Mountains basalt of Rollinson (1999) (solid circle) using partition coefficients quoted in Foley et al. (2002) and references therein. Following Foley et al. (2002), the residual assemblages are: 0.8 GPa, 40% amphibole, 45% plagioclase, and 15% orthopyroxene; 1.0 GPa, 32% amphibole, 15% plagioclase, 28% garnet, and 25% clinopyroxene; 1.6 GPa, 17% amphibole, 15% plagioclase, 15% garnet, and 52% clinopyroxene.

Generation of High Silica Continental Rocks: The Archean similar to the primitive mantle, and on this basis there appears to be little evidence for subduction or crustal contamination processes having contributed to their high thorium contents (Condie, 1994). Such evolved basalts are also less refractory, and so they are likely to yield the greatest amount of melt for a given temperature rise above the solidus in a basaltic pile undergoing anatexis. A second feature is that the TTG have negative Nb – Ta anomalies (Figure 5(a)), and these are widely regarded as indicative of subduction related magmatism. Figure 7(b) is a plot of Ta/La versus Ti/Zr, and TTG, like most granitic rocks, have both low Ta/La and Ti/Zr reflecting negative anomalies in both tantalum and titanium. A key feature of primitive oceanic arc lavas is that their negative Nb – Ta anomalies are not accompanied by relative depletions in titanium (Figure 2(b)), and thus low Ti/Zr ratios (Figure 7(b)). It follows that their low Ta/La values reflect subduction-related processes that have not fractionated Ti/Zr. In contrast, higher-K island arc suites such as that from the Aeolian Islands (Ellam et al., 1989), which typically have a greater contribution from subducted sediment, have low Ta/La and Ti/Zr values, as do most sediments (see the global subducting sediment (GLOSS) composition from Plank and Langmuir (1998)). Negative Nb – Ta and Ti anomalies are a feature of granitic suites (Figure 5), average crustal compositions, and the Archean TTG (Figure 7(b)). The key question is the extent that such negative Nb – Ta anomalies reflect the same processes as those responsible for the low Ti/Zr values, and whether low Ta/La ratios are formed by different processes in oceanic island arcs and granitic magmas. Clearly, if negative Nb –Ta anomalies are formed by different processes in high silica magmas, they cannot be regarded as a simple test of a subduction-related component in such magmas or in the generation of the bulk continental crust. The evidence from selected magmatic suites is not straightforward. TiO2 decreases with increasing SiO2 in most magmatic suites, but Ta/La fractionates differently. In the intraplate setting of Tenerife in the Canaries, for example, Ta/La increases from 0.09 to 0.13, with little change in Nb/Ta, as SiO2 increases from 48% to 60% in a basanite –phonolite suite modeled by closed system fractional crystallization (Ablay et al., 1998). In A-type granites Nb/La, and by implication Ta/La, increases in the most differentiated compositions (Figure 5(d)), but in many granitic rocks there is no simple intrasuite relationship between Ta/La and Ti/Zr. Most of the Idaho batholith samples plotted in Figure 7(b) have .60% SiO2, and the low

367

Ta/La and Ti/Zr values are a feature of the parental crustal melts (Clarke, 1990). Similarly, the TTG have low Ta/La and Ti/Zr and no clear intrasuite correlations between them, although that may, in part, reflect the large-scale sampling of many studies. An alternative approach is modeling on the basis of experimental data. In principle, low Ta/La and Ti/Zr values might reflect residual, or fractionating, titaniferous phases such as rutile, or amphibole. Rutile is widely invoked as a residual phase in slab melting models (see discussion in Kelemen et al. (2003)), but the recent experimental study of Foley et al. (2002) suggests that rutile incorporates Ta in preference to Nb and so coexisting melts will have high Nb/Ta. Within many TTG the data are scattered, but decreases in Ta/La are accompanied by increasing Nb/Ta. The TTG considered by Foley et al. (2002) had low Nb/Ta, and they argued that in absence of residual rutile, fractionation of titanium from zirconium must be accomplished primarily by amphibole. Figure 7(b) therefore compares the melts produced from an average LREE-enriched Archean tholeiite at different pressures, and thus with different residual assemblages, to natural, undifferentiated TTG, using the partition coefficients for low-Mg amphibole quoted by Foley et al. (2002). Although such modeling is sensitive to amphibole chemistry and the composition of the protolith, low-Mg amphibole can also impart the low Ta/La and Ti/Zr of the TTG. This is permissive evidence that the TTG may be formed by low degrees of melting (,10%) of fractionated basaltic rocks leaving a garnet amphibolite residue with little or no rutile, provided that the amphibole has Mg# values less than ,60. An amphibolebearing residue is also consistent with hornblende being a near-liquidus or fractionating phase in many TTG suites (Martin, 1995). An intrinsic advantage of the slab-melting model is that with continued subduction the voluminous garnet amphibolite residues (^ mafic cumulates) of TTG generation will transform into eclogite and ultimately founder into the deeper parts of the mantle, rather than being retained in the lower crust, as would be the case for derivation from a basaltic underplate. Such efficient “waste disposal” thus provides a ready explanation for the observation (see Chapter 3.01) that the bulk Archean crust appears to be, in general, more felsic than post-Archean crust (Rudnick, 1995). This is because most magmas of post-Archean subduction zones are derived from the mantle wedge rather than the descending oceanic slab (Martin, 1986; see below).

368 3.11.3.1.2

Granitic Perspectives Secular evolution in TTG composition

If fusion of subducted oceanic crust is the correct origin for many Archean TTG, we might anticipate some secular variation in TTG compositions as the Earth cooled. This is because the attendant decrease in the geothermal gradient along the Benioff plane should lead to an increasing depth of slab melting, and thus TTG generation. The angle of subduction is also likely to increase, since the envisaged decrease in the rate of seafloor spreading from the earliest Archean (Bickle, 1978) leads to an increase in plate dimensions and thus the average age of oceanic crust, and older, colder lithospheric slabs are less buoyant (Christensen, 1997). A direct implication is that the thickness of the overlying mantle wedge that slab-derived magmas must traverse, and potentially interact with, will increase with time, and this is consistent with observed changes in TTG compositions. In a compilation of TTG data, Martin and Moyen (2002) demonstrate that between 4.0 Ga and 2.5 Ga the least differentiated TTG evolve to progressively higher Mg# and Ni contents, consistent with greater assimilation of olivine during passage through the mantle for the younger TTG. In the same period strontium contents also increase appreciably, which is interpreted as evidence for a declining role for residual plagioclase, and thus an increase in the depth of TTG formation from the Early to Late Archean (Martin and Moyen, 2002). It appears that systematic changes in TTG compositions with time in the Archean can be related to likely changes in the conditions of melt generation, if the subducted slab model is adopted. Such secular chemical trends seem less easy to accommodate by models where TTG are formed by remelting basaltic rocks in the lower crust, given that the Mg# and Ni contents of mantlederived magmas (and thus the potential TTG source rocks) has decreased appreciably with time. Modern adakites extend to higher strontium and Mg# values than Archean TTG of similar silica contents (Smithies, 2000; Martin and Moyen, 2002). Higher strontium implies lower degrees of partial melting, compatible with a progressively cooling Earth, whereas the elevated Mg# suggests greater interaction with mantle peridotite (Martin and Moyen, 2002). The implication of the latter is that modern adakitic melts are generated even deeper than the TTG, near the threshold between slab melting and slab dehydration. Thermal conditions do not favor slab melting in most modern day subduction zones, except under the exceptional circumstances of anomalously

elevated heat flow associated with ridge subduction (Stern and Kilian, 1996), or the consumption of hot, newly created oceanic crust (Martin, 1986; Drummond and Defant, 1990; Peacock et al., 1994) (but see Chapter 3.18 for an alternative view). Hence, the volume of TTG-like rocks emplaced into the continental crust has waned dramatically since the Archean, and in some cases such compositions have been attributed to melt generation at depth within thickened continental arcs (see above). Nevertheless, the continental crust shares the low (subchondritic) Nb/Ta, high Zr/Sm signature relative to the primitive mantle that characterises the TTG (Foley et al., 2002), suggesting that this style of magmatism has made an important imprint on the bulk crustal composition. 3.11.3.1.3

A combined model for TTG formation

On the balance of the evidence discussed above, it seems likely that intraplate plumerelated magmatism and subduction-related magmatism were both important aspects of crust generation processes in the Archean (Albare`de, 1998). The key component of the TTG represents the second stage of crust generation, as partial melts of basaltic source rocks. It appears that those basalts are typically relatively evolved, in that they had high incompatible element contents (e.g., thorium; Figure 7(a)), and many greenstone belt volcanics have unfractionated Th/Ta suggesting that they were not generated by subduction-related processes, but may have had an intraplate or plume-related origin. It follows that in many cases the negative Nb – Ta anomalies of the TTG were not a feature of their basaltic source rocks, but were instead generated during partial melting of those basalts. Moreover, the low Ta/La ratios of the TTG appear to be linked to low Ti/Zr values, which characterizes most granitic rocks (Figure 7(b)). TTG were generated at depths where residual garnet, and not plagioclase, was stable but it is unclear whether sufficient crustal thickening on a large scale took place in the Archean to permit TTG generation at the base of the crust. Subduction ensures that TTG generation took place at sufficiently high pressures, and interaction of TTG magmas with peridotite in the mantle wedge offers an explanation for their high Ni and Mg#. However, it is curious why there are not more greenstone belt volcanic rocks with clear subduction signals, and the distinctive granite-greenstone belt outcrop patterns are not easy to reconcile with the linear sites of crust generation that characterize recent subduction zones. Models for TTG generation have been developed involving residual titaniferous phases

Generation of High Silica Continental Rocks: The Archean such as rutile, and low Mg# amphiboles, and a distinctive test may be that melts in equilibrium with residual rutile should have high Nb/Ta ratios (Foley et al., 2002). 3.11.3.2 Late Archean Granitic Rocks: The Onset of Major Intracrustal Melting Towards the end of the Archean, after ,3.0 Ga, greenstone belt volcanism and the TTG were largely supplanted by relatively potassium-rich granitic intrusives (monzogranite and granodiorite with subordinate tonalite), which comprise around 20% of presently exposed Archean cratons (Condie, 1993). As with Phanerozoic granites, “calc-alkaline” (mostly metaluminous to weakly peraluminous), strongly peraluminous and alkaline associations are recognized (Sylvester, 1995). According to Sylvester (1995), these are formed by remelting intermediate igneous and sedimentary compositions at mid- to lower-crustal depths, rather than wholly in the stability field of garnet as inferred for the TTG. The appearance of granitic plutons may therefore reflect stabilization of a critical volume and thickness of buoyant continental crust to permit prolonged weathering processes, formation of sedimentary basins, and ultimately intracrustal melting during collisional orogenesis (see below). 3.11.3.2.1

Generation of the Archean granitic types

Archean granitic rocks are broadly similar to their Phanerozoic counterparts, though some systematic differences can be resolved (Figure 5 and Table 3). In most cases these can be related to small differences in protolith composition and the melting zone lying at deeper crustal levels in the Archean. The Archean calc-alkaline association extends to higher Al2O3/TiO2, Na/K, Sr, Sr/Y, La/Yb, and Eu/Eup than Phanerozoic equivalents and these distinctive features were most likely inherited from a TTG component in the source region. The depleted HREE and yttrium of some Archean calc-alkaline granites requires residual garnet, suggesting formation in the lower crust, and this is much less common in Phanerozoic plutons. Average barium contents of Archean calc-alkaline granites are twice those of Phanerozoic plutons, and this is attributed to greater consumption of biotite during melting under higher thermal gradients in the Archean (Sylvester, 1995). However, high barium contents are not matched by higher rubidium, and it is therefore also possible that the elevated barium contents reflect a paucity of residual feldspar, consistent with a greater melting depth.

369

Interestingly, Phanerozoic calc-alkaline granites have distinctly higher Mg# than those of the Archean. This could reflect the presence of a “juvenile” mantle-derived component in the younger plutons, since, in contrast to the Archean, there is abundant field evidence for the involvement of such material in Phanerozoic granite genesis (see below). Strongly peraluminous Archean granitic rocks with relatively low CaO, Na2O, and Sr are most abundant in terranes that contain a major metasedimentary component, such as the Superior Province, Canada (e.g., Percival, 1989). As with their Phanerozoic equivalents, Archean peraluminous rocks are thought to derive largely by anatexis of metasediments (Feng and Kerrich, 1992; Sylvester, 1995). Like the Archean calc-alkaline plutons, the strongly peraluminous Archean plutons have, on average, higher Na/K, Sr, and Eu/Eup, but lower rubidium, scandium, yttrium, and HREE (Sylvester, 1995). These features directly mirror differences in composition between Archean and Phanerozoic graywackes (Taylor and McLennan, 1985, 1995), and therefore probably reflect secular variation in source composition, perhaps accentuated by a minor amount of residual garnet in the Archean. Archean alkaline plutons are similar in many respects to their Phanerozoic equivalents, with high SiO2 and rubidium, and prominent negative europium anomalies, though the characteristic high-field-strength element (HFSE) enrichment and CaO – Sr – Ba depletion (see Section 3.11.4.3.1) are less marked. Thorium contents are higher, and Ga/Al ratios are significantly lower in Archean alkaline rocks, the latter being a defining attribute of Paleozoic alkaline plutons (Collins et al., 1982). Given uncertainties with pinpointing the sources of alkaline plutons of Phanerozoic age (Section 3.11.4.3.2) these differences are difficult to evaluate petrogenetically. However, the lower Ga/Al and higher thorium may signify derivation from precursors that were less refractory than those postulated for Phanerozoic alkaline granites, and the higher CaO – Sr– Ba suggest that feldspar fractionation is less important in the evolution of Archean alkaline plutons. On a more general note, the average Archean calc-alkaline and alkaline granitic plutons in the compilation of Sylvester (1995) have consistently higher Th/U ratios than Phanerozoic granites (Figure 8), even though this is not a property of the estimated bulk Archean crust (see Condie (1993) and Taylor and McLennan (1995) for compositions of the Archean crust). Ellam et al. (1990) also noted that the time-integrated Th/U ratios of Archean granites calculated from their lead isotopes were much higher than those in

370

Granitic Perspectives strongly peraluminous Archean granitic rocks, which were sourced from such protoliths. The higher Th/U ratios of the calc-alkaline and alkaline Archean granites may partly reflect lack of a sedimentary ingredient, in tandem with a contribution from relatively thoriumenriched mantle-derived magmas and/or crustal fluids during melting.

14

Archean granites 12

Th/U (ppm)

10 8 6 4

Bulk crust

2 0

Calc-alkaline

Strongly peraluminous

Alkaline

Figure 8 Plot comparing the average and range of measured Th/U ratios in the various Archean granitic types (shaded squares) with those of post-Archean granites of the same chemical group (open squares). All data are from Sylvester (1995), though the range of Th/U ratios indicated for the post-Archean granites incorporates data from the LFB (sources as for Figure 6).

average crustal compositions. However, the strongly peraluminous Archean granites appear to have slightly lower average Th/U ratios than their Phanerozoic counterparts. Besides the possibility that this feature is not a product of weathering and/or alteration, there are two potential explanations for these observations. First, it may be that aqueous fluids were not as important in the Archean for the generation of the calc-alkaline and alkaline granite types as in the Phanerozoic. Alternatively, uranium may not have been in the appropriate (oxidized) valence state to be readily mobilized in low fO2 Archean fluids (in contrast to U4þ, U6þ is highly soluble in aqueous fluids). This second point is consistent with other lines of evidence. For example, to reconcile variations in lead isotope data shown by MORB, numerical simulations by Elliott et al. (1999) require that the upper mantle evolves towards markedly lower Th/U ratios after ,2.2 Ga. These authors attribute the decreasing Th/U ratios to a change in the oxidation state, and thus enhanced mobility of uranium in the more oxygen-rich environment at the surface of the post-Archean Earth, such that this element is preferentially recycled into the mantle as a component of altered oceanic crust (see also McCulloch and Bennett, 1994). As a result of the reduced mobility of uranium during continental weathering in the Archean, shales of this period have lower mean Th/U ratios than their younger counterparts (McLennan and Taylor, 1980) and this feature was therefore acquired by

3.11.3.2.2

Tectonic implications

By analogy with “classical” Phanerozoic orogens, the geological context of Archean granites, and the temporal trend of some suites from deformed, calc-alkaline plutons to undeformed alkaline rocks, suggests formation in a collisional to post-collisional tectonic setting (Sylvester, 1995 and references therein). Should this prove correct, it supports the presence of rigid plates in the Late Archean and the operation of conventional plate tectonics by this time. Nevertheless, analogues of the “Cordilleran” type continental arc batholiths, belts of calcic tonalite-dominated suites associated with abundant coeval mafic rocks, are apparently absent in the Archean, and these are one of the hallmarks of convergent plate margins. This point can be construed as either (i) further evidence against subduction-related models for the generation of Archean continental crust, or (ii) reflecting differences in the thermal regime, such that heat input from mantle magmas was less essential in triggering granite genesis in the hotter Archean crust than in young orogens (Section 3.11.4). Typical calc-alkaline igneous suites formed above contemporary subduction zones are thought to have evolved from basaltic parental magmas derived from the mantle wedge, and it has been suggested that this may have been less common in the Late Archean if magmas were largely derived by partial melting of the downgoing slab (Martin, 1986). The compositional parallels between Archean granitic rocks and their Proterozoic to Phanerozoic counterparts suggest that the style of crustal differentiation was similar during these periods. The implications for crustal evolution are therefore discussed in the next section.

3.11.4 GENERATION OF HIGH SILICA CONTINENTAL ROCKS: II. PROTEROZOIC TO PHANEROZOIC GRANITIC ASSOCIATIONS The sedimentary record reveals a shift in upper crustal composition near the Archean – Proterozoic boundary (Section 3.11.2.3),

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 371 suggesting a concomitant change in the degree, and perhaps style of crustal differentiation, and possibly in the composition of new crust. Magmatically, this change is signaled by the emplacement of large volumes of relatively potassic, incompatible-element-rich granitic plutons from about the Late Archean onwards, and the localization of the TTG style of magmatism to specific convergent plate boundary settings (Martin, 1986). Most younger plutons are distinguished from the Archean TTG by less REE fractionation and by having more conspicuous negative europium anomalies across the entire compositional range (Figure 6), translating this character to post-Archean sediments. Notably, Proterozoic and younger granitic suites are typically not bimodal and include less differentiated, more mafic granitic compositions. Although there are provincial differences between some Proterozoic granitic suites and those in the Phanerozoic, no systematic differences in average global sedimentary compositions or near-surface heat flow can be resolved for these time periods (Taylor and McLennan (1985, 1995); see also the compilations of Condie (1993)). Small fluctuations may be masked by sedimentary averaging and recycling processes, but the constancy of sedimentary compositions implies that the processes of crustal generation and differentiation were similar throughout the postArchean, and that net crustal additions from the Archean onwards were of approximately the same composition.

3.11.4.1 The Importance of Intracrustal Melting, and the Mantle Connection Since potassic granitic magmatism is usually inferred to be the primary agent of crustal differentiation in Proterozoic and Phanerozoic terranes, a number of issues need to be addressed: (i) the efficiency of elemental fractionation during the generation and evolution of these granitic magmas; (ii) the extent that these can be reconciled with estimates of the gross chemical structure of the continental crust; and (iii) the amount and composition of new crust formed during such granitic episodes. Several lines of evidence indicate that mid- to shallowlevel intracrustal melting became significantly more important after the Archean. First, in contrast to the TTG, the near ubiquity of negative europium anomalies and flat HREE patterns indicate that magma generation occurred at relatively shallow, mid-crustal depths with plagioclase, rather than garnet or amphibole, as a residual or fractionating phase.

In some cases, the negative europium anomaly may also have been inherited from the source rocks, but this still requires that these have experienced an intracrustal differentiation cycle involving plagioclase, or (in the case of sediments) are derived from such rocks. Second, large discrepancies between crystallization ages and neodymium model ages (e.g., McCulloch and Chappell, 1982) confirm that older (crustal) material was present within the source regions of granitic magmas. Third, there is commonly a component of inherited zircon in post-Archean plutons (e.g., Paterson et al., 1992; Williams et al., 1992), again consistent with a source dominated by recycled crustal materials. Moreover, as phase relationships confirm that intracrustal melting is fundamentally dependent on the availability of solidus-lowering volatiles, post-Archean crustal evolution has been summarized as “no water, no granite; no oceans, no continents” (Campbell and Taylor, 1983). The water required for melting can occur as a free (supercritical) vapor phase (“fluid-present melting”), or be liberated by the peritectic decomposition of hydrous minerals, such as micas and amphiboles, which occurs at higher temperatures and is referred to as fluid-absent or “dehydration melting” (Thompson, 1982). These scenarios have vastly different geochemical consequences for the derivative melts (Section 3.11.4.2). In nature, crustal anatexis is likely to be a progression from fluid-present to fluid-absent conditions as aH2O decreases during melting, especially where the incipient fluid-rich melts are extracted from the source region. Although intracrustal melting characterizes the post-Archean Earth, the emplacement of granitic rocks in some oceanic settings remote from continental influences (Section 3.11.4.3) suggests that old crustal rocks are not a prerequisite for silicic magma formation, and that juvenile mantlederived materials may be important. Moreover, even in continental settings, there is generally some evidence for mantle involvement, either as syn-plutonic basaltic dikes, coeval gabbros, swarms of mafic enclaves, or primitive isotopic signatures. A synergy between intracrustal melting and mantle-derived magmatism in granitic genesis is to be expected, since intrusion of hot, basaltic magma is likely to induce crustal fusion (Huppert and Sparks, 1988), and thus granites may incorporate variable proportions of mantle- and crust-derived components. Thus we envisage three different scenarios for granite generation, involving pure crustal sources, pure mantle sources, and variable combinations of these. Each scenario, and its relevance for post-Archean crustal evolution, is explored in the following sections by reference to specific case studies from well-documented Phanerozoic terranes.

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Figure 9 Field photographs of the selected granite types (by T. Kemp). (a) Migmatite outcrop from the CambroOrdovician Glenelg River Complex (southeastern Australia) showing the deformation-enhanced segregation and accumulation of locally derived partial melts into a dilatant shear zone. Such segregation between partial melt and restite epitomises mid-crustal anatectic terranes (Brown, 1994). In this case, the leucogranite in the shear zone is compositionally equivalent to the kilometer-scale leucogranitic plutons of the area, and supports the viability of a genetic nexus between in situ migmatites and granitic bodies. (b) Quenched gabbroic diorite globules hosted by a subvolcanic alkali-feldspar granite on Jersey, Channel Islands (UK). Note the incorporation of alkali feldspar phenocrysts from the granite into the mafic pillow on the bottom left, suggesting that the granitic host was partly solidified prior to injection of the mafic magma. (c) Mingling between a syn-plutonic basaltic dike and the Tuross Head Tonalite, Moruya Batholith, eastern LFB (location indicated on Figure 15). The basaltic dikes and pillows have chilled margins and a tholeiitic to mildly alkaline geochemistry. As with the host tonalite, they have high Sr

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 373 3.11.4.2 Peraluminous Leucogranites—Granites Derived from Old Crustal Protoliths

3.11.4.2.1

Conditions of leucogranite formation

Strongly peraluminous leucogranites (granites or granodiorites with ,5 vol.% of mafic minerals; Le Maitre (1989)) with metasedimentary protoliths arguably provide the least ambiguous examples of granites generated exclusively from remelting of pre-existing crustal rocks (Patin˜o Douce, 1999). They represent the large-scale vertical redistribution of mass within the crust, rather than net growth of the continental crust. These rocks typically occur in continental collisional zones, such as in the Himalayas (Section 3.11.4.2.2), North American Appalachians (e.g., Pressley and Brown, 1999), and the Hercynian fold belts of western Europe (e.g., Strong and Hanmer, 1981; Wickham, 1987). The pluton-forming processes are directly observable in outcrop, typically due to the uplift and exhumation of deeper crustal levels, permitting geochemical and isotopic linkage between protolith and leucogranite. The viability of the process is confirmed by experimental studies, where fusion of the metasedimentary (usually metapelitic) protolith in the laboratory at the relevant P – T-aH 2 O conditions consistently reproduces the felsic, strongly peraluminous compositions of the leucogranite (e.g., Patin˜o Douce and Harris, 1998; Castro et al., 2000). One implication of this is that leucogranites approximate pure anatectic liquid compositions, which is confirmed by field relationships (below).

Detailed field and geochemical studies have established that leucogranite emplacement represents the culmination of complexly interacting, multistage melt segregation and extraction processes, commonly controlled by syn-anatectic deformation (see Figure 9(a)), in which fractional crystallization, localized crystal accumulation, and separation of entrained refractory phases have important roles (see Brown, 1994; Solar and Brown, 2001). Isotopic data further indicate that rather than being simply derived from a compositionally homogeneous source, leucogranitic bodies comprise a poorly blended amalgamation of partial melt batches extracted from a range of metasedimentary protoliths (e.g., Deniel et al., 1987; Krogstad and Walker, 1996; Pressley and Brown, 1999; Kemp, 2001). However, there is rarely evidence for the direct involvement of coeval mantle-sourced magmas in leucogranite genesis (an exception is documented by Kemp (2003)). In the absence of heat transfer from mafic liquids (that is, assuming only conductive heat flux from the mantle), modeling suggests that temperatures sufficient for fluid-present partial melting (, 650 –700 8C) may be achieved in the lower part of the crust by internal radioactive heating and thermal relaxation associated with crustal thickening, even using average continental geotherms (England and Thompson, 1986). However, the same calculations reveal that the higher

contents (to 800 ppm) and exhibit primitive isotopic compositions, ranging to 1Nd ¼ þ8 (Keay et al., 1997). These basalts are used as the primitive mantle-derived end-member in the mixing models of Keay et al. (1997) and Collins (1998). However, the high Na, high Sr signature is only present within granitic suites of the easternmost Bega Batholith. (d) Injection and mingling between dioritic and granitic magmas in the eastern Bega Batholith, LFB. Compared to the basalts in Figure 9(c), the dioritic magmas are compositionally evolved and exhibit certain textural features, notably hornblende-pyroxene-rimmed quartz ocelli and replacement of mafic minerals by biotite, that suggest they had experienced significant hybridization prior to mingling with the granite. (e) Swarms of hornblende-bearing “microgranular” enclaves in a hornblende granite of the eastern Bega Batholith. Note the morphological and textural diversity of the enclaves and the general lack of crenulate (quenched) margins, as is typical of microgranular enclaves in high-level granitic plutons. (f) Close-up view of a microgranular enclave in a hornblende tonalite of the Jindabyne suite, LFB. Enclaves such as this have essentially basaltic compositions, with 49– 51% SiO2, though the enclave population as a whole exhibits the complete spectrum through to intermediate compositions. (g) Array of enclave types in the cordierite-bearing Cowra Granodiorite, LFB. Metasedimentary enclaves range from discrete tabular objects (center left) to more diffuse, partially assimilated bodies (top, near lens cap). Note the irregular microgranular enclave (left) and the darker margin on the spherical microgranular enclave on the center right. The rim is biotite-rich and thought to result from the diffusional influx of K2O and H2O from the host magma, promoting biotite crystallization at the expense of pyroxene; relict hypersthene is sometimes preserved in the center of the enclave (Vernon, 1990; Maas et al., 1997). The numerous clots of ferromagnesian minerals imparting the heterogeneous textural aspect are typical. Many of the clots result from the physical disaggregation of metasedimentary and microgranular enclaves, arrested examples of which are conspicuous in outcrops. Other clots, such as those consisting of large cordierite – sillimanite – hercynite ^ garnet lumps may be the refractory residue of partial melting. (h) Closer view of a microgranular enclave contained by a lower silica cordierite granite of the Kosciuszko Batholith. The massive structure is typical. Although many enclaves of this type have a superficially “basaltic” appearance, their mineralogy usually mimics that of the host, such that they are rich in biotite and commonly contain cordierite and aluminosilicate.

374

Granitic Perspectives

Figure 9

temperatures required for fluid-absent, mica dehydration melting reactions (.700 8C) are only attained under anomalous conditions of perturbed (i.e., steeper) geotherms and/or enhanced radiogenic heat flux (England and Thompson, 1986; Thompson and Connelly, 1995). As muscovite (but less so biotite, see Patin˜o Douce and Harris (1998), figure 3)) dehydration melting equilibria have shallow, positive dP=dT slopes (Thompson, 1982), much larger melt proportions are generated during adiabatic or near-isothermal decompression, such as occurs during exhumation by normal faulting (England and Thompson, 1984, 1986).

(continued ).

This could partly explain the observation of Castro et al. (2000) that peraluminous leucogranites of collisional belts tend to be associated with extensional, rather than compressional structures. Muscovite dehydration melting during decompression at mid-crustal depths has been inferred for leucogranite generation in the Iberian massif, Spain (Castro et al., 2000), and the Himalayan orogen (Harris and Massey, 1994; Searle et al., 1997; Section 3.11.4.2.2). However, high precision monazite dating by Harrison et al. (1999) reveals that the Manaslu Intrusive Complex of the central Himalaya was emplaced during two pulses of muscovite

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 375

Figure 9

dehydration melting separated by ,4 Myr, and to achieve this by decompression alone requires unrealistically high denudation rates. On the basis of numerical simulations, Harrison et al. (1997, 1999) instead propose that the diachronous leucogranitic magmatism was triggered by shear heating along major Himalayan fault systems. Nabelek and Liu (1999) also conclude that mica dehydration melting leading to Proterozoic leucogranite magmatism in the Black Hills (South Dakota, USA) was promoted by shear heating accompanying tectonic unroofing.

(continued ).

Irrespective of whether crustal anatexis is driven by thermal relaxation attending the collapse of tectonically thickened orogens or frictional heating along active faults, the important point is that leucogranite generation via muscovite-involved melting can be achieved without heat advection from mantle-derived magmas. The thermo-tectonic models of England and Thompson (1984, 1986) also predict that P –T paths might intersect higher temperature melting reactions involving breakdown of biotite in metapelites and hornblende in

376

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Figure 9

amphibolites during crustal thickening where the initial geotherms are abnormally “hot” (see also Thompson, 1999). More mafic (biotite-rich) granitic magmas that contain abundant peritectic reaction products of biotite dehydration, such as cordierite, garnet and/or orthopyroxene, and even metaluminous granitic plutons in the case of amphibolite fusion, could result from this mechanism. However, two additional points stem from this. First, field examples of intracrustal melting in Proterozoic to Phanerozoic metamorphic belts consistently involve metasedimentary or, more rarely, intermediate igneous

(continued ).

precursors. This remains true even in the deeper crustal exposures, such as the Ivrea Zone of Northern Italy, where metapelitic rocks experienced granulite-facies anatexis close to the Moho (Voshage et al., 1990). Melting of meta-basic protoliths is rare, exceptions being exhumed subduction complexes where a ready fluid supply is available (e.g., Sørensen, 1988). This is unlike in the Archean, where extensive migmatitic metabasite terranes occur, such as in the Superior Province of the Canadian Shield (e.g., Sawyer, 1991). This presumably reflects the higher solidus temperatures of meta-basic

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 377 rocks (on average ,822 8C in fluid absent conditions; see Petford and Gallagher (2001)), which were more easily attained in the higher temperature conditions of the Archean. Second, as noted by several authors (e.g., Clemens and Vielzeuf, 1987), the strongly modified geotherms required for biotite dehydration melting in the middle crust are most easily reconciled by an enhanced heat input from the mantle, either by upwelling of asthenosphere beneath attenuated crust, delamination of mantle lithosphere, or direct incursion of hot, mantlederived liquids into the crust (underplating or intraplating; see Thompson (1999) for an evaluation of these possibilities). The latter is clearly the case for fusion of metapelites in the Ivrea Zone, which occurred at the roof of a crystallizing basaltic sill (Voshage et al., 1990), and variations of this scenario have been extensively modeled (e.g., Huppert and Sparks, 1988; Bergantz and Dawes, 1994; Petford and Gallagher, 2001). Moreover, in the dynamic environment of multiple pulses of basaltic injection into partially melted crust it seems implausible that the derivative anatectic melts should remain physically and chemically isolated from their heat sources. Indeed, geochemical, isotopic, and experimental evidence, and the content of distinctively igneous-textured microgranular enclaves, attest to a mantle-derived component in many peraluminous granitic plutons generated by biotite dehydration of metasedimentary protoliths, such as the cordierite-rich Hercynian granodiorites of western Europe (Barbarin, 1996; Di Vincenzo et al., 1996; Castro et al., 1999; this aspect is addressed in Section 3.11.4.4). It is for this reason that such plutons are not as suitable as leucogranites for modeling the effects of closed-system intracrustal differentiation.

3.11.4.2.2

Low-strontium leucogranites

The majority of strongly peraluminous leucogranitic plutons documented in the literature contain strontium contents that are less than 200 ppm, and these are hereafter referred to as low-strontium leucogranites. Perhaps the most classic and thoroughly studied example of such rocks are the Miocene leucogranites of the Himalayan orogen (e.g., Le Fort, 1981; Searle and Fryer, 1986; Le Fort et al., 1987; Castelli and Lombardo, 1988; Crawford and Windley, 1990; Scaillet et al., 1990; Inger and Harris, 1993; Searle et al., 1997; Harrison et al., 1999 and references therein). These comprise laccolithic, polyphase plutons associated with numerous dike and sill complexes that crop out

intermittently along the 2,000 km strike length of the Himalayan range, and they are emplaced into a partly migmatitic metasedimentary sequence. The size of individual laccolithic plutons varies, but ranges up to 3,000 km3 for the Manaslu Complex (Guillot and Le Fort, 1995). Although geochemically evolved (i.e., high silica, potassium, and rubidium contents), the leucogranites are considered to approximate primary minimum melts that are unmodified by fractional crystallization (Inger and Harris, 1993) and disengaged from the source without significant entrainment of refractory phases. Traceelement modeling (Harris and Inger, 1992; Harris et al., 1995), isotopic constraints (Deniel et al., 1987; France-Lanord and Le Fort, 1988; Inger and Harris, 1993), and experimental studies (Scaillet et al., 1995; Patin˜o Douce and Harris, 1998) suggest that the partial melts that coalesced into the plutons were generated by muscovite dehydration melting of adjacent metapelitic schists at 600 – 800 MPa and 700 – 750 8C. Biotite – muscovite leucogranites are thought to result from a larger melt fraction formed at higher temperatures (, 700 –750 8C) than the tourmaline-bearing leucogranites (, 700 8C) (Scaillet et al., 1995; Ayres et al., 1997). The average biotite and tourmaline leucogranites from the Langtang Valley of the Nepalese Himalaya (Inger and Harris, 1993) are plotted in Figure 10(a), compared to a putative source rock (kyanite schist). Selected trace-element ratios are shown in Figure 11. During partial melting leading to Himalayan leucogranite formation, plagioclase was a significant residual phase, in addition to biotite, peritectic alkali feldspar, and accessory minerals (Harris and Inger, 1992; Harris et al., 1995). The melts are therefore calcium poor and are characterized by depleted strontium and barium, leading to higher Rb/Sr and Rb/Ba, and lower Eu/Eup compared to their inferred metasedimentary protoliths (Figures 10(a) and 11). The Himalayan leucogranites are also depleted in elements compatible in the other residual phases, such as biotite (titanium, niobium) and accessory zircon (zirconium, hafnium, HREE), monazite (LREE, thorium, uranium), and apatite (phosphorus, yttrium), reflecting the minimal dissolution of these minerals during anatexis. This is a consistent feature of leucogranites worldwide, manifested by their low mafic mineral content. Even lower yttrium and ytterbium contents in some samples reflect the presence of residual garnet (Harris et al., 1995). Most Himalayan leucogranites reached chemical equilibrium with restitic zircon prior to melt extraction, as shown by

378

Granitic Perspectives

Himalayan leucogranites 10

1

0.1

biotite granite Tm granite kyanite schist

Bulk crust normalized Ba

U

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(a)

GRC leucogranites 10

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residue schist Ba

Bulk crust normalized

U

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Sm

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0.01 Rb

Th

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(b)

Figure 10 Bulk crust normalized multi-element diagrams comparing the chemical fractionation between source, melt and residue for (a) Himalayan leucogranites (Tm, tourmaline) from the Langtang region, Nepal (data from Inger and Harris, 1993) and that for (b) an average leucogranite from the Glenelg River Complex (GRC) in southeastern Australia; the field of in situ migmatite leucosomes is shaded (from Kemp, 2001).

the concordance between their zirconium contents and those predicted by zircon solubility equations (Ayres et al., 1997; Figure 12) However, some leucogranites, including most tourmaline-bearing varieties, contain very low LREE concentrations that imply incomplete equilibration with residual monazite (Figure 12). This probably reflects the marginally lower temperatures of tourmaline granites, and thus more sluggish diffusion rates of the REE, and the short residence time of the melt in the source (Ayres et al., 1997, and see below).

3.11.4.2.3

High-strontium leucogranites

A group of peraluminous leucogranite plutons with low K/Na ratios and unusually elevated strontium contents (250 –450 ppm) occur in the Glenelg River Complex of southeastern Australia (Kemp and Gray, 1999; Kemp et al., 2002). These contrast strongly with the predominant low-strontium leucogranite variety discussed above (see Table 3) and thus warrant further examination here. The Glenelg River Complex comprises part of the Cambro – Ordovician Delamerian Orogen

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 379 100

Cordierite granite Himalayan schists Himalayan LG GRC schists GRC LG

LFB turbidites

1 Sr/Nd

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Figure 11 Trace-element ratio plots emphasizing the different chemical effects of fluid-present and dehydration melting during leucogranite generation from metasedimentary protoliths. The LFB cordierite granites (Kemp, 2003, unpublished data), and the field of the Lachlan Ordovician turbidites (B. W. Chappell, unpublished data) are shown for comparison. (a) Rb/Sr versus Eu/Eup, (b) Sr/Nd versus Rb/Ba, and (c) Eu/Eup versus Sr/Eup. The compositional vectors resulting from igneous and sedimentary processes are indicated in (c). Some of the high-grade Glenelg River Complex (GRC) schists plot along the igneous fractionation trend, which cannot reflect weathering but is consistent with extraction of a partial melt component, as seen in the leucogranites of the area. Note the contrasting trends defined by the Himalayan (dehydration melts) and GRC (fluid-present melts) leucogranites, despite the source rocks being similar in composition (see also Figure 10). Also note the limited overlap between the LFB cordierite granites and the Ordovician sedimentary rocks, their potential protoliths.

(equated with the trans-Antarctic Ross Orogen, Stump et al. (1986)), which is thought to have formed by arc-continent collision in the Late Cambrian (Kemp, 2003). During the crustal thickening attending collision, a range of quartzofeldspathic to semi-pelitic metasedimentary rocks (metagraywackes) experienced low temperature anatexis at ,650 – 680 8C and 500 – 800 MPa, during which biotite was essentially refractory. The low temperatures, coupled with the persistence of prograde muscovite and paucity of potassium – feldspar – sillimanite assemblages, suggest that partial melting occurred in the presence of excess aqueous fluid, rather than by decomposition of hydrous minerals (Kemp and Gray, 1999). This situation is considered

anomalous, given the limited porosity envisaged for metamorphic rocks in the middle crust, and the high solubility of water in silicic melts at these pressures (,10 wt% at 500 MPa (Johannes and Holtz, 1996)) (see discussion in Clemens and Vielzeuf, 1987). Fluid present melting has been advocated in the Trois Seigneurs Massif of the French Pyrenees, though here the derivative leucogranites have low (,100 ppm) strontium contents (Wickham, 1987). A high-strontium (331 ppm) “trondhjemitic” leucogranite intrusion in the Himalayas is attributed to fluid-fluxed melting of a metapelitic source at the onset of anatexis (Patin˜o Douce and Harris, 1998). In the Glenelg River Complex, melting of the fertile metagraywackes may have been triggered by

380

Granitic Perspectives cordierite granites Himalayas GRC leucogranites

Zr / Zr*

1

0.1 0.1

1 LREE / LREE*

Figure 12 Plot of Zr (ppm measured in the rock)/Zrp (Zr contents in the melt at zircon saturation) versus LREE (La to Gd, excluding Eu, measured)/LREEp (LREE contents in the melt at monazite saturation) to assess and compare the roles of zircon and monazite melting in water-present and dehydration melts (solubility equations of Watson and Harrison (1983) for zircon, and Montel (1993) for monazite). The concentrations of Zr and LREE in the melts are largely buffered by the stability of zircon and monazite during anatexis (Ayres et al., 1997). The “wet” Glenelg River Complex (GRC) melts have very low Zr and LREE concentrations and are therefore undersaturated in zircon and monazite, reflecting the minimal solution of these minerals during low temperature melting (, 700 8C), and perhaps some occlusion by residual micas. In contrast, the Zr and LREE contents of the LFB cordierite granites exceed the predicted solubility of zircon and monazite. The cordierite granites are thus oversaturated in these minerals at the inferred crystallization temperature (750 8C, Chappell et al., 2000) and melt water content (5 wt%, from Clemens and Wall, 1981). This is consistent with the physical entrainment of refractory accessory phases by these magmas, which results in unusually high concentrations of those elements that normally have low solubility in granitic melts (Chappell et al., 1987; Chappell, 1996b).

ingress of fluid from an external source, such as dehydration of underthrust lower-grade schists, or cooling of fluid-rich arc magmas at depth. The muscovite-bearing leucogranitic bodies of the Glenelg River Complex are embedded within the migmatitic metasedimentary sequence, where they have lenticular or sheet-like geometry and may be up to 5 km across (Kemp and Gray, 1999). They have interleaved, or transitional boundaries with the migmatites, and can be traced into veins and patches of in situ leucosome via networks of dikes and sills (Figure 9(a)). These field relationships suggest that the leucogranites represent ponded accumulations of partial melts derived from the adjacent metasedimentary sequence. Strong support for this is that the leucogranites are isotopically identical to the quartzofeldspathic migmatites (Kemp and Gray, unpublished data) and geochemically equivalent to the centimeter-scale in situ leucosomes hosted by these rocks (Figure 10(b)). Furthermore, due to the efficient, deformation-enhanced segregation and extraction of the constituent partial melts from their biotite-rich residues, it is likely that leucogranites approach liquid compositions. An average Glenelg River Complex leucogranite is plotted in Figure 10(b), compared to

an inferred protolith and complementary residue of partial melting. The distinctively high strontium and barium, but lower Rb/Sr and Rb/Ba, in the leucogranite relative to the protolith (Figure 11) reflect the paucity of residual plagioclase, the stability of which during melting is depressed relative to the micas by excess fluid (Patin˜o Douce and Harris, 1998). The enhanced consumption of plagioclase during anatexis in the muscovite stability field is therefore responsible for the higher strontium contents of the Glenelg River Complex leucogranites relative to most Himalayan leucogranites (Figures 6(a) and 10(b)). The Glenelg River Complex leucogranites are also depleted in titanium, rubidium, HFSE, and the REE, which are sequestered by residual biotite and accessory minerals. However, unlike many Himalayan leucogranites, both the zirconium and LREE contents of the highstrontium leucogranites are lower than those expected from theoretical zircon and monazite solubility equations (Figure 12). This indicates that in the Glenelg River Complex chemical equilibrium between melt and refractory accessory phases was not attained prior to melt extraction. This could reflect a combination of: (i) the low temperatures of anatexis impeding

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 381 diffusion of zirconium and the LREE, (ii) occlusion of accessory minerals from the partial melt by enclosure within residual micas, this being actually observed in migmatite melanosomes, and (iii) short melt residence times, such that the rate of melt extraction from the source region exceeded that required for equilibration with dissolving zircon and monazite in the residue. The last of these accords with the field evidence that partial melt segregation and migration are driven by syn-anatectic deformation (see Figure 9(a)) and are thus likely to have been rapid (Harris et al., 2000). Modeling based on monazite dissolution rates (assuming magmatic temperatures of 700 8C) suggests that the tourmaline-bearing Himalayan leucogranites migrated from their sources in less than 104 yr (Ayres et al., 1997). Using the higher magmatic temperatures inferred from experimental studies (750 8C), this reduces to ,100 yr (Harris et al., 2000), since under these conditions the leucogranites are also undersaturated in zircon, and zirconium homogenization rates in a melt are three orders of magnitude faster than monazite (Ayres et al., 1997). Although there is large inherent uncertainty in these calculations (mainly in melt water contents and diffusion rates for zirconium and the LREE), similarly short melt residence times are apposite for the Glenelg River Complex leucogranites, given their undersaturation in zircon and monazite. The anomalously low zirconium contents (,10 ppm) in two leucogranites (which are not otherwise chemically unusual) could be due to heterogeneous distribution of zircon grains in the analyzed sample, or that zircon crystals in the protolith were mostly armored by refractory minerals. Another curious aspect of the high-strontium leucogranites in the Glenelg River Complex is their positive europium anomalies, since most leucogranites, and indeed Phanerozoic granitic rocks in general, have pronounced negative europium anomalies (Figures 11(a) and (c)). Large positive europium anomalies are commonly exhibited by small volume migmatite leucosomes generated under a variety of P – T-aH2O regimes, where they are variously attributed to near-complete extraction of the feldspathic component of the source during high degrees of melting (Watt and Harley, 1993; Johannes et al., 1995), disequilibrium melting of feldspar (Watt et al., 1996), localized plagioclase accumulation (Sawyer, 1987; Milord et al., 2001), or diffusional equilibration between leucosome and melanosome (residue) during crystallisation (Nabelek, 1999) or protracted cooling (Fourcade et al., 1992). However, positive europium anomalies in kilometer-sized leucogranitic plutons require special circumstances.

Diffusional exchange with the residue can be ruled out, since the leucogranitic melts were isolated from melanosomes during crystallisation, and in any case plagioclase was virtually exhausted from the residual assemblage. Textural evidence and the near cotectic compositions are not consistent with the leucogranites being feldspar cumulates (Kemp, 2001), though minor feldspar accumulation cannot be precluded. Instead, the positive europium anomalies are attributed to extensive disequilibrium melting of plagioclase, followed by rapid extraction of the partial melt from the source migmatites. Under these circumstances, europium is preferentially released into the melt and is therefore decoupled from the trivalent REE, bound in the refractory accessory phases that are concentrated in the residue. A final aspect of this style of leucogranite generation is that the most important radiogenic heat-producing elements (thorium, uranium) are enriched in the residues of melting (i.e., apatite, zircon, and monazite), and thus retained in the deeper crust. This contrasts with the situation modeled by Sandiford et al. (2002), in which these elements behave incompatibly during melting and are transported to the upper crust, resulting in long term cooling, and thus stabilization, of the granitic source terrane. However, in the case of leucogranite generation the lower crust remains hot following the evacuation of partial melt, and may even undergo second stage melting, depending on the ambient geotherm and the size of the thermal pulse induced by radioactive decay. The residues from the original leucogranite formation are hydrous (mica-bearing) and thus could yield significant amounts of granitic melt during such an event, via higher temperature (,850 8C) biotite-dehydration melting (Patin˜o Douce and Johnston, 1991). Apart from postdating leucogranite emplacement, these melts would be recognized by elevated K/Na ratios, and would have much higher Rb/Sr, but lower Eu/Eup than the high-strontium leucogranites, in these respects more closely resembling the upper continental crust. This second-stage melting scenario may explain the intrusion of peraluminous magmas in some terranes that postdate the deformational peak. 3.11.4.2.4

Relevance for crustal evolution

The effective separation of the silicic liquid from the metasedimentary source that epitomizes leucogranite generation maximizes the geochemical signature of intracrustal melting (Sawyer, 1998), and yields compositions that are strongly fractionated relative to the bulk continental crust (Figure 5(c)). Operating alone, the extreme chemical fractionation associated with this style

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of magmatism would tend to drive the upper crust towards a haplogranitic composition, that is quite different from the currently estimated upper crustal composition. In particular, muscovite dehydration melting of pelitic rocks under the conditions of Himalayan leucogranite genesis, while resulting in upwards concentration in lead, rubidium, potassium, and uranium, cannot produce the enrichment in Sr, Ba, REE, Nb – Ta, Zr – Hf that typifies the upper crust relative to the inferred bulk crust composition (Figure 5(c)). This problem is most acute for strontium and barium since these soluble elements will be preferentially removed by seawater and thus their concentrations will be further diluted in the upper continental crust from those observed in pristine igneous rocks (see Section 3.11.6). Fluid-present minimum melts can contribute towards enrichment in Ba and strontium in the upper continental crust, though the low Rb/Sr and positive europium anomaly of these melts contrasts with the characteristic high Rb/Sr and large negative europium anomaly of this reservoir. We therefore conclude that, although important for enriching the upper crust in the most incompatible elements, intracrustal melting leading to leucogranite generation cannot be the principal agent driving differentiation of the continental crust. This is not unexpected, given that leucogranites are typically volumetrically minor in the context of most orogenic belts, especially those developed at convergent plate margins. In fact, the majority of granitic rocks exposed in the upper crust are considerably more mafic than leucogranites (compare to other Phanerozoic granites in Table 3), suggesting the operation of additional processes and/or derivation from different sources under different anatectic conditions.

3.11.4.3 Granites Formed from Juvenile Mantle-derived Materials These granites can be generated by extreme differentiation of basaltic magmas, or by remelting mafic mantle-derived materials that have had a short residence time in the crust. Both models effectively result in the formation of new crust and in differentiation, but as the resultant granitic magmas are isotopically primitive, it may be difficult to isolate which of the two processes was responsible for their formation. The production of broadly granitic magmas from juvenile, mantle-derived materials is most easily demonstrated in oceanic settings remote from continental influences, for example, (i) the development of the volumetrically minor “plagiogranites” of mid-ocean ridge systems, now

preserved as part of ophiolite complexes (e.g., Pedersen and Malpas, 1984); (ii) alkaline granites of the Kerguelen plateau in the Indian Ocean (Giret, 1990); (iii) Icelandic dacites and rhyolites (Furman et al., 1992; Jonasson et al., 1992); and (iv) the primitive “M (mantle)-type” quartz diorites of intraoceanic island arcs, such as those of New Britain (Whalen, 1985) or the Aleutians (Perfit et al., 1980). The last group is potentially of relevance to the continental crust for two reasons. First, they comprise early intrusions in some continental margin arc-related batholiths, such as the Peninsula Ranges Batholith, California (Silver and Chappell, 1988), suggesting that they represent the initial stages of the formation of thick continental crust above subduction zones. Second, they are considered to represent the source rocks for metaluminous granites generated in subsequent thermal events unrelated to subduction (e.g., Chappell and Stephens, 1988). Nevertheless, at the same silica content, oceanic arc granites are most unlike the upper continental crust in having lower rubidium, thorium, uranium, LREE, and markedly depleted niobium (Figure 5(b)). Continental arc granites have higher LILE contents, though niobium and phosphorus remain lower than upper and bulk crustal values. Based on analogy with arc andesites (e.g., those of the Andes, see above), it is likely that some of these features result from crustal contamination, rather than being an intrinsic property of the original magma, and thus their direct relevance for the compositional evolution of the bulk crust is not immediately obvious (see Section 3.11.6 and Figure 21).

3.11.4.3.1

A-type granites

Distinctively silicic, potassium-rich granitic rocks that sometimes have mantle-like isotopic signatures also occur in intracontinental rift zones and postcollisional environments where they usually (but not invariably, see Whalen et al., 1987; King et al., 1997), postdate convergent deformation. They tend to be emplaced at shallow crustal levels, commonly occurring as subvolcanic ring complexes associated with crystal-poor rhyolites (see Bonin, 1986). Such plutons are collectively referred to as “A-type” (Loiselle and Wones, 1979), to reflect their commonly alkaline nature, the content of anhydrous mafic minerals, and their emplacement in intraplate or “anorogenic” settings. Since the original definition by Loiselle and Wones (1979), a large literature has accumulated on A-types, and it has become increasingly apparent that rocks of this designation are

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 383 lithologically and compositionally diverse. For example, they may be peralkaline, and contain sodium-rich mafic minerals, such as the aegirinearfvedsonite granites of Nigeria (Kinnaird et al., 1985), or they can be weakly peraluminous to metaluminous, containing hornblende or ferrohastingsite, like the Siluro-Devonian A-types of the LFB, eastern Australia (see review by King et al., 1997). Both types may even occur in the same area (e.g., Corsica, Poitrasson et al., 1995). Nevertheless, some common traits are recognized. Many A-type rocks are hypersolvus granites (Tuttle and Bowen, 1958), that contain either a single alkali feldspar, or an unmixed (perthitic) potassic feldspar mantled by albite (“Rapakivi” texture), this being in accord with low-pressure crystallization. A-types typically contain high temperature anhydrous phases, such as pyroxene and fayalite (or their relicts), and late-crystallizing (interstitial) biotite and amphibole, both of which may be fluorine-rich (e.g., Anderson, 1983; Wormald and Price, 1988). Fluorite is a common accessory phase, and zircons in samples studied by ion microprobe mostly lack older, pre-magmatic cores (King et al., 1997). These features suggest formation from high temperature (.900 8C), completely molten magmas with low to moderate water content, an interpretation that is consistent with the shallow emplacement levels and has been confirmed experimentally (Clemens et al., 1986). Many A-type granites contain primary ilmenite and thus have fO2 values below the fayalite – magnetite – quartz (FMQ) buffer (Loiselle and Wones, 1979; Anderson, 1983; Turner et al., 1992a; C. D. Frost and B. R. Frost, 1997), in contrast to the relatively oxidized, magnetite-bearing metaluminous granites that are characteristic of arcs, which typically lie 1 –3 log units above FMQ (Ishihara, 1979; Frost and Lindsley, 1991). Chemically, most A-types are characterized by high K 2O, Fe/(Fe þ Mg), and Rb/Sr, and enriched to extreme high field strength element (HFSE, especially zirconium, niobium, and yttrium) and REE3þ concentrations, with strongly depleted calcium, barium, strontium, and europium contents (see summary in King et al., 2001). These features implicate either extended fractional crystallization or small degrees of partial melting to generate the incompatible element enrichment (Section 3.11.4.3.2). The large negative europium anomalies also reflect the low fO2 and associated high Eu2þ/Eu3þ ratio (see discussion of the effects of fO2 on europium partitioning in Section 3.11.5). A-types are also distinguished by elevated Zn and Ga/Al ratios, and the latter has been taken as diagnostic (Collins et al., 1982; Whalen et al., 1987). However, although the chemical integrity

of A-types is marked at the lowest silica contents, towards high silica there is compositional convergence with fractionated metaluminous granitic suites that otherwise lack A-type character (Figure 13(a); see King et al., 1997). Chemical discriminants are most effective when applied to associations of rocks, where the less evolved members are also represented (King et al., 1997). Alternatively, given the difficulties with simple Harker variation diagrams, reasonable separation between A-type and non-A-type granites can be achieved using more complex chemical parameters that emphasize the HFSE enrichment and silica-rich, calcium-poor nature of the former (Figure 13(b)). Another curious feature of A-types is their uneven distribution throughout the geological record. Rocks of this nature are encountered in Archean cratons (Section 3.11.3.2.1), though they are far subordinate to the TTG and other granitic types. Similarly, A-type magmas are volumetrically minor in Phanerozoic orogenic belts; they comprise only 0.6% of the vast granitic batholiths of the LFB (Chappell et al., 1991), discussed in Section 3.11.4.4. However, A-type rocks are considerably more abundant in the Proterozoic, e.g., the vast Proterozoic anorogenic granitic provinces of the southwestern US, Adirondack Mountains and eastern Canada (Anderson, 1983), and the Rapakivi-textured granitic sheets of Greenland (Brown et al., 1992). Large volumes of anorogenic, HFSE-enriched granitic (Wyborn et al., 1992) and volcanic (Creaser and White, 1991) units were emplaced across Australian Proterozoic terranes in the 1.8 –1.5 Ga interval. Similarly, Rapakivi-textured A-type batholiths intruded the Early Proterozoic crust of southern Finland between 1.65 Ga and 1.55 Ga at the close of the Svecofennian orogeny (Haapala and Ramo, 1992). In all of these areas, silicic A-type rocks are invariably coeval with mafic intrusions, and this is also a feature of some younger A-types (see below), hinting that a mantle input is intrinsic to their genesis. Interestingly, the largest volumes of anorogenic granites in the Australian Proterozoic and Greenland were emplaced immediately following the 1.9 Ga “super-event” of Condie (1998, 2000; see Figure 2), where massive crustal growth was thought to be triggered by mantle plumes. In keeping with this, some chemical features of the A-type granites, especially the HFSE-enrichment, are reminiscent of plumerelated basalts. However, most A-type granites (except the most fractionated varieties), and their coeval mafic rocks, have negative niobium anomalies (i.e., low HFSE/LREE). These are

384

Granitic Perspectives 700 hornblende granites 600

A-type granites

Zr (ppm)

500

400

300

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100

0 55

60

65 70 SiO2 (wt.%)

(a)

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Zr (ppm) / M

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0

0 (b)

50

100

150

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Nb+Y (ppm)

Figure 13 (a) Variation of Zr (ppm) as a function of SiO2 (weight percent) showing the compositional overlap between the LFB hornblende granites and A-type granites in the most siliceous rocks (data sources as for Figure 6). The steep trend defined by A-type granites is compatible with zircon saturation in felsic rocks of the group (King et al., 2001). Note the comparatively flat trend of the hornblende granites and the downwards inflection at ,74% SiO2. (b) Plot of Zr (ppm)/M versus Nb þ Y, where the contrasting granitic types show minimal overlap. The parameter “M” is the atomic ratio [Na þ K þ (2·Ca)]/(Si·Al) (from Watson and Harrison, 1983) which tends to be higher in A-type granites than hornblende granites of similar silica content, reflecting the Ca-deficiency of the former. Thus, the Zr/M ratio is directly proportional to the zircon saturation temperature of Watson and Harrison (1983), although this is only rigorously applicable to melts just saturated in zircon, and invalid for peralkaline granites.

not a feature of intraplate basalts, but the low Nb/La ratios of A-type granites are coupled with low Ti/Zr (Figure 5(e)) consistent with fractionation by a titaniferous phase. Some support for

a spatial link with mantle plumes is that silicic rocks with A-type affinity are recorded from hotspot settings, including continental (e.g., Yellowstone, Hildreth et al. (1991); East African rift,

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 385 MacDonald et al. (1987); Etendeka province, Namibia, Schmitt et al. (2000)) and oceanic, (e.g., Kerguelen, Giret (1990); Ascension Island, Harris (1983)) environments.

3.11.4.3.2

Petrogenetic models for A-type granites

Reflecting their diversity and distinctive compositions, a number of petrogenetic schemes have been proposed for the origins of A-type suites. These essentially fall into two categories, involving crustal and mantle sources, respectively. In either case, fractional crystallization is considered to be the dominant process controlling the geochemical evolution of A-type granite suites, consistent with the inference that magmas were completely liquid at high temperatures, and the extreme degrees of differentiation recorded by the geochemistry. For models advocating a crustal source, the protolith composition, and the conditions of magma generation are somewhat unusual. These models, underpinned by trace element and experimental evidence, include low degrees of partial melting of a dry, granulitic residue depleted by the prior extraction of granitic melt (Collins et al., 1982; Clemens et al., 1986; Whalen et al., 1987), remelting metaluminous tonalites or granodiorites (Anderson, 1983; Creaser et al., 1991; Anderson and Morrison, 1992; Skjerlie and Johnston, 1993; Patin˜o Douce, 1997; King et al., 1997, 2001), and fusion of a dehydrated mafic to intermediate “charnockitic” lower crust (Landenberger and Collins, 1996); very low water activities and fO2, and high temperatures (.900 8C) are intrinsic to all models. The first two of these scenarios have shortcomings (see reviews by Turner et al. (1992a) and C. D. Frost and B. R. Frost (1997)) and there is little reason that any should be universally applicable. The elevated temperatures necessary to generate A-type liquids from refractory or low H2O crustal sources undoubtedly requires major heat input from mantle-derived magmas (e.g., Landenberger and Collins, 1996), and so many A-types may incorporate a large juvenile component in their lineage. Indeed, physical mingling and hybridization between A-type magmas and coeval basaltic intrusions has been documented (e.g., the Newfoundland Appalachians, Whalen and Currie (1984); Figure 9(b)), but whether mantle-derived magmas have made a material contribution to other “crustally derived” A-types, is difficult to assess and relies heavily on isotopic evidence. Haapala and Ramo (1992) demonstrate that the neodymium and lead isotopic composition

of the Finnish Rapakivi granites parallels the evolution of the 1.9 Ga Svecofennian crust, which is consistent with the granites representing anatectic melts of these older rocks. Conversely, the Proterozoic Rapakivi granites of Greenland have low initial 87Sr/86Sr and positive epsilon neodymium values, suggesting that juvenile crust and/or mantle magmas were more appropriate sources (Brown et al., 1992). Isotopic evidence seems less conclusive in other cases where A-type granites have 1Nd and initial 87 Sr/86Sr ratios that straddle crustal and mantle values (Poitrasson et al., 1995) For example, King et al. (1997) show that the Siluro –Devonian A-type plutons from the LFB have a large spread in epsilon neodymium values (23 to þ5) at the time of crystallization that overlap the range shown by the least evolved hornblende granites of the area. This has been taken to indicate source heterogeneity and it appears to preclude a significant sedimentary component within these rocks (King et al., 1997), as is true of A-types in general (see the compilation of Turner et al., 1992a). However, some Lachlan A-types have higher 1Nd and lower initial 87Sr/86Sr ratios than the isotopic envelope defined by hornblende granites; we shall return to this point in Section 3.11.4.3.3. In contrast, some A-type granites, including most peralkaline varieties, are ascribed a largely mantle derivation (e.g., Loiselle and Wones, 1979; Bonin, 1986; Bedard, 1990; Turner et al., 1992; Whalen et al., 1996; C. D. Frost and B. R. Frost, 1997; Han et al., 1997; Jahn et al., 2000; Schmitt et al., 2000), and these are the most pertinent in the present context of crustal growth. The basis for many of these models are primitive isotopic compositions (Javoy and Weiss, 1987; Turner et al., 1992a; Whalen et al., 1996; Han et al., 1997; Jahn et al., 2000; see Section 3.11.4.3.3), with trends to more evolved values signifying assimilation or mixing with older crustal materials (Foland and Allen, 1991; Kerr and Fryer, 1993; Poitrasson et al., 1995; Schmitt et al., 2000). Given the distinctive traceelement chemistry of A-type granites, a somewhat “enriched” (intraplate) mafic precursor may be required, and this is variably attributed to small degrees of mantle melting during lithospheric extension (Barbarin, 1999; Jahn et al., 2000), melting lithospheric mantle fertilized by prior subduction (Turner et al., 1992a,b; Whalen et al., 1996), or derivation from a plume source (Schmitt et al., 2000).

3.11.4.3.3

A-type case studies

An example of Phanerozoic crustal growth during A-type magmatism is discussed by

386

Granitic Perspectives

Turner et al. (1992a) with reference to the silicic magmas of the bimodal padthaway suite in southern Australia. These magmas were emplaced during postcollisional extension immediately following the convergent deformation of the ca. 500 Ma Delamerian orogeny. The silicic rocks, occurring as shallow level plutons and rhyolitic volcanics, are distinguished mineralogically by Rapakivi textures, interstitial annitic biotite, and sporadic fayalite and Pidgeonite, with accessory fluorite and large euhedral zircons. Pyroxene-olivine assemblages and coexisting Fe– Ti oxides preserve equilibria established at 900 –1,000 8C and ,100 MPa, and constrain an oxygen fugacity that is 2– 3 log units below the FMQ buffer (Turner et al., 1992a). By analogy with the experimental study of Clemens et al. (1986), crystallisation proceeded under strongly water-undersaturated conditions, with final H2O contents below 3% (Turner et al., 1992a). Geochemically, the rocks are highly differentiated and exhibit pronounced A-type affinity, contrasting with the older metaluminous and peraluminous plutons of the same area (Foden et al., 1990; Turner et al., 1992a). In particular, initial isotopic ratios are relatively primitive (1Nd , þ1 to 22.7, 87Sr/86Sr , 0.706) and indistinguishable from associated tholeiitic basalt and gabbro intrusions (Turner et al., 1992a), with which they are intimately intermingled (Turner and Foden, 1996). Accordingly, the silicic A-type magmas are considered to have evolved from these contemporaneous basaltic magmas by protracted (,90%) fractional crystallization, and thus exemplify the addition of new granitic crust. The characteristic iron enrichment trends, low fO2 and high incompatible element concentrations are all explained by this process (Turner et al., 1992a). Importantly, the enriched incompatible element concentrations of the coeval basaltic rocks, particularly high La/Yb and low Nb/La, coupled with positive epsilon neodymium values, suggests derivation primarily from enriched domains within the subcontinental mantle lithosphere, rather than the asthenosphere (Turner, 1996). The postcollisional magmatism has, therefore, been attributed to delamination or convective thinning of the overthickened lithosphere following tectonic convergence, allowing upwelling of the hot asthenosphere, and facilitating contact melting of the overlying enriched mantle (Turner et al., 1992b; Turner, 1996). A lithospheric mantle derivation has also been invoked for the parental magmas of other postcollisional granitic suites (e.g., the Adamello Massif; Blundy and Sparks, 1992). A variation of this scenario, proposed for the Proterozoic Rapakivi granites of North America

by C. D. Frost and B. R. Frost (1997), is that A-types form by remelting young, underplated tholeiitic basalts, including their evolved, ferrodiorite differentiates, in areas of crustal extension or hot spot activity. This model has the advantage of explaining the bimodal character of A-type suites, though inevitably the hot, dry silicic melts would undergo fractional crystallization upon ascent into the shallow crust, amplifying their distinctive compositional features. Hence, there is conceptual overlap with the model of Turner et al. (1992a) and the implications for crustal growth are unchanged. The high-temperature A-type magmas formed in this way are very susceptible to crustal contamination, especially since they have low strontium contents and 87Sr/86Sr ratios, which may partly explain the characteristic dispersion of isotopic compositions exhibited by these rocks (Turner et al., 1992a; C. D. Frost and B. R. Frost (1997)). Although the basaltic progenitors in the C. D. Frost and B. R. Frost (1997) model are depleted tholeiites derived from upwelling asthenosphere, in principle more enriched and isotopically evolved mantle sources could also be involved. In view of this, it is striking that the “aluminous” A-type granites from the LFB fall within the isotopic range of the lithosphere-derived mafic magmas documented by Turner (1996), and thus potentially could be sourced from such magmas rather than pre-existing crust. One sample of King et al. (1997) has an 1Nd value of þ5 at the time of crystallization, which was clearly inherited from a mantle source, or a juvenile magma derived therefrom. Some syn-plutonic basaltic rocks are associated with A-type granites of the Wangrah suite (King et al., 2001), though unfortunately their isotopic compositions and some key trace-element concentrations (e.g. niobium) are not presented. However, these basalts have enriched zirconium (,250 ppm) and cerium (to 89 ppm) contents and thus could easily yield A-type-like liquids upon high-temperature partial melting, since zirconium and cerium are unlikely to be buffered by residual zircon or allanite in the mafic protolith. 3.11.4.3.4

Relevance for crustal evolution

The average Padthaway suite A-type pluton of Turner et al. (1992a), together with metaluminous and peralkaline A-type granite samples from the LFB are plotted in Figure 5(d). The markedly enriched HFSE and REE contents are salient features, and of all granitic types considered, only these magmas can contribute significantly towards the concentration of Nb – Ta and Zr – Hf, and depletion in TiO2 in

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 387 Australia. Although lithologically diverse, these define an apparently simple, overlapping array on the 1Nd versus initial 87Sr/86Sr diagram (Figure 14), which has been almost universally used to infer a large scale mixing process between primitive, depleted mantle-like magmas and evolved crustal end-members (e.g., Gray, 1984; Faure, 1990). Yet, such models may not reconcile issues such as why samples with differing mineralogy and major element compositions should have similar trace-element ratios, and how mixing alone can explain large chemical differences between granitic rocks that are isotopically similar. It is also unclear whether the granites represent liquid compositions or contain unmelted refractory material from the source (“restite”), and evolved by the differential “unmixing” of such material White and Chappell, 1977). This has been a particular debate for the Lachlan granites and it has shaped much of the wider debate over granitic generation (e.g., Vernon, 1983; Wall et al., 1987; Chappell et al., 1987; Collins, 1998, 1999 versus Chappell et al., 1999, Chappell et al., 2000).

Lachlan Fold Belt

DM 5 10% 20% 0 Central Vic.

e Nd

the upper crust. Despite the HFSE enrichment, negative Nb – Ta anomalies are present even in isotopically primitive (mantle-like) A-types, and this also mirrors the continental crust (the high Nb/La ratio in the peralkaline rock possibly reflects the late-stage fractionation of REE-rich accessory phases). On the basis of this, and despite their limited areal extent, it is tempting to speculate that A-type magmatism may have made a more important contribution to the HFSE budget and HFSE/LREE ratio of the upper crust than has hitherto been appreciated. However, elevated Zr – Hf and Nb – Ta concentrations are linked to enhanced solubility of these elements in the alkaline melt via high magmatic temperatures, extensive dissolution of accessory phases in the source (King et al., 1997), and the formation of Na – F complexes (Collins et al., 1982), all of which require anomalous circumstances during granitic generation. In the case of a purely mantle heritage, the slightly positive 1Nd values suggest moderate-term enrichment of the source in the LREE. Furthermore, A-type magmas have pronounced depletions in barium, strontium, and phosphorus, and though these are less well developed in the metaluminous A-type in Figure 5(d), such depletions are not typical of the upper continental crust. In order to assess the broader significance of A-type magmatism for crustal evolution, future studies need to resolve several outstanding issues. These include (i) ascertaining whether partial melting of young mafic protoliths or fractionation from mantle magmas is the most important for the generation of these plutons; (ii) where enriched mantle sources are invoked, constraining the conditions leading to the formation of these; and especially (iii) why A-type magmas are so voluminous in the Proterozoic, as compared to the Archean or Phanerozoic periods.

–5

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3.11.4.4 Granites with Juvenile Mantle and Crustal Sources: The Lachlan Case Study 3.11.4.4.1

Mixing or unmixing?

Granitic intrusions of exclusively and unambiguously mantle provenance are rare in Phanerozoic orogenic belts, and in many cases petrological evidence demands derivation of these magmas from mixed sources, which included both mantle-derived materials and older crustal protoliths. However, constraining the nature and proportions of these components has often proved formidable. Such complexities are now explored in a case study of granitic rocks from the Paleozoic Lachlan Fold Belt (LFB) in southeastern

Figure 14 Initial Nd– Sr isotope relations in the LFB granites (data sources are: McCulloch and Chappell, 1982; Chappell et al., 1991; McCulloch and Woodhead, 1993; Elburg, 1995 b,c; Maas et al., 1997; Keay et al., 1997; Anderson, 1998; Waight et al., 2000; Kemp, unpublished data, and unpublished data quoted by Chen et al., 1990). Cordierite granites from the central Victorian portion of the LFB plot away from the main array and appear to comprise a separate petrogenetic system, as is consistent with other geochemical criteria such as anomalously high Ba contents (Anderson, 1998). The hyperbola represents a putative mixing line between the average Ordovician turbidite (av. sed.) and the depleted mantle end-member (DM) determined by McCulloch and Chappell (1982), and encapsulates most hornblende and cordierite granites (crosses represent 10% mixing intervals).

388 3.11.4.4.2

Granitic Perspectives Two contrasting granite types

Any examination of granites within the LFB is indebted to the pioneering and systematic studies of Chappell and White and co-workers (see Chappell and White, 2001), through which large field, petrographic and geochemical data sets are available and form the basis for petrogenetic models. However, despite intensive research, roughly since the early 1970s, there remain lingering questions and controversial aspects concerning the formation of these granitic rocks. The purpose of this case study is to focus on some of these specific problems in the context of crustal growth and evolution. The LFB is a 700 km wide segment of a 3,600 km long orogenic system that developed along the eastern Gondwana margin from the Early Ordovician to Devonian (Figure 15). The tectonic setting of the belt remains contentious, although the complex patterns of deformation, magmatism, and metamorphism seem to require multiple, migrating subduction zones (Gray, 1997; Soesoo and Nicholls,1999). The LFB has two main components, a monotonous succession of mature (quartz- and clay-rich) Ordovician turbidites, and an extraordinary volume of granitic rocks. The turbidites apparently accumulated on an oceanic substrate and were subject to episodic deformation, low-grade regional metamorphism and massive igneous intrusion from ,450 Ma to 340 Ma (Gray and Foster, 1997). Granites and spatially related volcanic units comprise approximately one quarter of the exposed Paleozoic geology of the orogen; gabbros or other relatively mafic rocks are rare. Excluding the volumetrically minor A-types, and following the seminal suggestion of Chappell and White (1974), the granites can be subdivided into two lithological categories, each of which outcrop in approximately equal proportions. The first group consists of metaluminous to weakly peraluminous tonalite, granodiorite and adamellite (ASI , 1.1) that are sometimes mingled with coeval basaltic intrusions (Figures 9(c) and (d)) and contain hornblende-rich, igneous-textured microgranular enclaves (Figures 9(e) and (f)). As these rocks have precipitated hornblende over much of their compositional range, they are here referred to as “hornblende granites.” Such plutons are predominant in the easternmost part of the LFB, where they form vast, meridionally trending batholiths (Figure 15). In contrast, the second group comprises strongly peraluminous granodiorites and adamellites (ASI . 1.0) that mostly contain cordierite, sometimes accompanied by sillimanite, garnet or

muscovite, and for convenience are termed cordierite granites. Cordierite granites only crop out some distance inboard of the continental margin (Figure 15), where they occur in composite batholiths with hornblende granites. In the Kosciuszko Batholith they are consistently older on the basis of intrusive relationships (White et al., 1976), though radiometric ages of the two granite types overlap in the Berridale Batholith further east (Williams et al., 1975). The most distinctive field attribute of cordierite granites is that they contain a diverse array of metasedimentary enclaves and mica-rich clots, which become increasingly numerous in the lowest silica samples (Figure 9(g)). These enclaves usually have higher metamorphic grade and more complex structural histories than the host turbidites (Fleming, 1996) and there are migmatitic examples with refractory, melt-depleted compositions (e.g., Chen et al., 1989; Anderson, 1998; Maas et al., 1997). Most metasedimentary enclaves were therefore entrained from a zone of deep crustal anatexis, and may therefore be utilized as structural and compositional “windows” to the unexposed mid-crust (Fleming, 1996; Anderson et al., 1996). Other enclaves include diopside-bearing “calc-silicate” varieties that perhaps represent fragments of metamorphosed marl (Chen et al., 1989). Less abundant are unstructured “microgranular” enclaves which are darker and finer-grained than their hosts (Figure 9(h)). The mineralogy of these generally mimics that of the host, though orthopyroxene is common and some enclaves of this type contain actinolitic amphibole (Vernon, 1990; Maas et al., 1997). The origin of the microgranular enclaves continues to be controversial (e.g., Maas et al., 1997 versus White et al., 1999) and in some ways is the key to understanding the formation of the host granite. The complex textures are most easily reconciled by igneous crystallization, implying that the microgranular enclaves are intermingled, variably hybridized globules of a coeval, more mafic magma (Vernon, 1983, 1984, 1990). Nonetheless, field evidence for such interaction involving the cordierite granites has yet to be documented. Furthermore, the microgranular enclaves in LFB cordierite granites exhibit an array of lithological and chemical variability, and it seems inevitable that they have multiple origins. The distinction between the LFB cordierite granites and strongly peraluminous leucogranites is also worth noting. Although both have the same ASI range, cordierite granites extend to considerably more mafic (cordierite- and biotiterich) compositions (Table 3). At the felsic end of the compositional range the differences are less marked, though cordierite granites retain systematically higher niobium, thorium, REE, titanium,

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 389 35˚ NT WA

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Figure 15 Simplified geological map of southeastern Australia (location in inset), showing the distribution of cordierite (light gray) and hornblende (black and dark gray) granites and equivalent volcanic suites within the eastern part of the LFB, and the location of the specific granitic batholiths and plutons (CG, Cooma Granodiorite) referred to in the text (modified after Chappell, 1996a). The volumetrically minor A-type granites, as well as small gabbro bodies, are grouped with the hornblende granites.

and yttrium abundances than the leucogranites (Figure 5), and show systematic enrichment in zirconium and the LREE compared to the concentrations predicted by zircon and

monazite solubility equations (Figure 12). The latter is compatible with the abundance of older, pre-magmatic zircon cores in the LFB cordierite granites, as documented from various plutons by

390

Granitic Perspectives 1.6

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Figure 16 (a) (a1 – a14) Chemical variation diagrams for the LFB hornblende and cordierite granites, and microgranular (MME) and metasedimentary enclaves (MSE) contained by cordierite granites. The field of the Ordovician turbidites is enclosed by a thin dashed line, and the Cooma Granodiorite is circled by an unbroken line. Note that the metasedimentary enclaves with very high CaO contents (.10 wt%) are calc-silicates. (b) (b1, b2) Plots of ASI versus SiO2 (top) and total FeO (bottom). The compositional trends generated from a granitic minimum melt (“MM”) by restite entrainment/increasing degrees of melting (thick arrow), contamination by Ordovician turbidites, magma mixing and fractional crystallization (thin dashed arrows) are indicated for the cordierite granites. Scattering of cordierite granites into the Cooma Granodiorite field suggests that partially melted (diatexitic) Ordovician turbidites were assimilated by ascending granitic magmas (see also Collins, 1996). Note the similar trends shown by the cordierite granites on the ASI versus silica and K2O/Na2O versus silica plots. That these correlate with initial 87Sr/87Sr suggests that the dominant compositional control was source-related for these plutons, specifically reflecting different degrees of weathering of the various sedimentary protoliths (see also discussion of Figure 19). Data sources are; granites, as for Figure 6; Microgranular enclaves, White et al. (1977),

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Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 391

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Chen et al. (1989), Elburg and Nicholls (1995), Elburg (1996b,c), Maas et al. (1997), Anderson (1998), Kemp (2003), unpublished); Metasedimentary enclaves; White et al. (1977), Chen et al. (1989), Maas et al. (1997), Anderson (1998), Kemp (2003), unpublished); Ordovician turbidites, Wyborn and Chappell (1983), Munksgaard (1988), Chappell et al. (1991), Anderson (1998), B. W. Chappell (unpublished data), Kemp (2003), unpublished data).

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Granitic Perspectives 2.2 Ordovician turbidites

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Williams (1992), Williams (1995), Anderson et al. (1996), Elburg (1996a), and Keay et al. (1999). 3.11.4.4.3

0

Geochemical and isotopic comparisons

In accordance with the mineralogical differences, the LFB cordierite and hornblende granites have different major element compositions, which are shown in an overall sense by Figure 16. The cordierite granites have a relatively restricted silica range (mostly . 65%), and towards the lower end they exhibit pronounced clustering on variation diagrams. This contrasts with hornblende granites, which show smooth and continuous compositional variation back to basaltic silica values. More importantly, cordierite granites define tighter chemical arrays that at silica values of ,65% project towards higher TiO2 FeOt, MgO, Cr, Pb, Rb, K/Na, Rb/Sr, and ASI, and lower Na2O, CaO, and Sr than most hornblende granites. The ASI is perhaps the best discriminant, as the

(continued ).

trends of each granite type diverge and differences are largely maintained over the full compositional range. Differences occur between specific samples on mantle normalized diagrams, where a cordierite granite can have higher contents of most trace elements, and lower strontium than a typical hornblende granites of similar silica (Figure 17). Nevertheless, the patterns and hence the trace-element ratios are strikingly similar, which is perhaps surprising given the ASI difference, and the widespread application of the I- and S-type nomenclature. Chemical contrasts are less apparent at higher silica, where the respective trends defined by the two granite types converge. The exception is P2O5, which shows a marked increase at the higher silica contents in cordierite granites, in contrast to a strong decrease for felsic hornblende granites, reflecting differences in the behavior of apatite (Chappell, 1999; see below). It is also important to note that although the geochemical arrays of both groups are approximately linear, at silica contents above ,74% there is an exponential

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 393 10

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increase and decrease in elements incompatible (rubidium) and compatible (strontium, barium) in feldspars respectively. Similar inflections are evident for zirconium and the LREE, and coincide with a sharp increase in ASI in the most differentiated cordierite granites. Cordierite granites extend to more evolved Sr –Nd (and O; O’Neil and Chappell, 1977) isotopic compositions than the hornblende granites. Those evolved compositions overlap with those of the Ordovician turbidites (Figure 14), and yield older, Proterozoic neodymium model ages. It follows that the cordierite granites were derived from source regions that have higher average Rb/Sr ratios than those of the hornblende granites (see Section 3.11.5.1 and Figure 19). Such features are consistent with cordierite granites containing a greater proportion of recycled crust, and this accords with the abundant inherited zircon cores contained by these rocks. Older, pre-magmatic zircons are also recognized in some hornblende granites (Williams et al., 1988, 1992) and their microgranular enclaves (Chen and Williams, 1990). However, cordierite granites extend to more primitive isotopic values than the metasedimentary basement and overlap the range shown by the most evolved hornblende granites. As a result, both granite types define a wellpublicized 1 Nd – initial 87Sr/86Sr hyperbola,

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Figure 17 Bulk crust-normalized multi-element diagram comparing a typical cordierite granite of the LFB (the Jillamatong Granodiorite of the Kosciuszko Batholith) and two hornblende granites at , 67% SiO2. The hornblende granites are from the eastern (Cobargo Granodiorite) and western (Glenbog Granodiorite) sides of the Bega Batholith, respectively (all data from Kemp, 2003, unpublished). The similar shape of the cordierite granite pattern to that of the hornblende granite from the western Bega Batholith is striking, and cannot simply be attributed to compositional convergence at high silica contents. The hornblende granite from the eastern Bega Batholith has less of the signature attributed to a sediment component, notably in higher Sr and Eu, but lower Pb and Rb/Ba.

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Figure 18 Plot of initial 87Sr/86Sr versus silica for (a) LFB hornblende granites (data sources as for Figure 14) and (b) cordierite granites (modified from figure 1 in Chappell et al., 1999), with annotated vectors summarizing the potential petrogenetic processes operating to produce the pattern of isotopic variation. The alternative is that the variation was inherited from a heterogeneous source (see text for discussion). Note the similarity between this figure and the ASI – SiO 2 plot in Figure 16b1. Tie lines in (a) join granitic rocks of the Moruya Suite, whose nearly closed-system isotopic variation could be caused either by fractional crystallization (Collins, 1996, Collins, 1998) or restite unmixing (Griffin et al., 1978), both of which will impart dispersion parallel to the silica axis. There is a vague general trend towards decreasing initial 87Sr/86Sr with silica for the hornblende granites, the possible petrogenetic significance of which awaits more detailed study within individual hornblende granite suites.

bracketed between the depleted mantle and Ordovician sediments (Figure 14). However, less well recognized (but see McCulloch and Chappell, 1982) is that there is no simple or systematic relationship between the position of specific rocks on the isotopic array and their whole rock composition. For the hornblende granites, this could, to some extent, reflect the modifying effects of fractional crystallization (Figure 18(a)) Nevertheless, a curious feature is that many of the most mafic hornblende granites are also the most isotopically evolved, such as some plutons from the Bega Batholith (Blue Gum Tonalite, 62.8% SiO2, 1Nd 2 8.9) and

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Berridale Batholith (Finister Tonalite, 64% SiO2, 1Nd 2 8) (data sources in Figure 14). The pattern of isotopic variation shown by the cordierite granites is even more complex (Figure 18(b)). Many felsic cordierite granites have lower 87Sr/86Sr than the mafic samples, and within an individual cordierite granite suite initial 87Sr/86Sr increases with silica to , 0.715 at 68 – 69% SiO2 before decreasing steadily to ,0.709 at 74% SiO2. Such observations impose strict limitations on magma mixing models proposed for the LFB granites.

3.11.4.4.4

Constraints on the origin of the LFB granites

Although the Sr– Nd isotopic array suggests that the LFB cordierite and hornblende granites formed simply by variable degrees of mixing between shared source components (depleted mantle-like magma, and an evolved crustal material), this is precluded by the different geochemical trends defined by the two granitic types. For this reason, it has been persistently argued that the hornblende and cordierite granites were generated separately from contrasting meta-igneous and metasedimentary source rocks, respectively (Chappell and White, 1974; White and Chappell, 1977; Chappell and White, 2001 and references therein). However, it is unclear why rocks with very different ASI, which supposedly reflect the disparate protoliths, should have such similar minor and trace-element patterns, as evident in Figure 17. Thus, we envisage that mantle-derived magmas and older crustal rocks were common ingredients in the formation of both the cordierite and hornblende granites, but that these materials were incorporated differently, and in different proportions. As will be discussed below, we also consider that the petrogenesis of each granitic type was coupled during Lachlan orogenesis. Cordierite granites. The cordierite granites have the greatest crustal affinity, though in some ways their petrogenesis and magmatic evolution is the most problematic. In a general sense, the petrography and geochemistry of these plutons accords with derivation from aluminous metasedimentary protoliths (Chappell and White, 1974). Such a process is evident in several small metamorphic complexes in the central and eastern LFB, where the Ordovician turbidites locally attain migmatite grade and have gradational contacts with small cordierite-rich, diatexite-like plutons. The most thoroughly studied of these, the Cooma Granodiorite, is geochemically and isotopically equivalent to its migmatitic envelope, and demonstrably formed by a closed-system bulk

“mobilization” of partially melted quartzofeldspathic metasedimentary rocks (Pidgeon and Compston, 1965; White et al., 1974; Flood and Vernon, 1978; Munksgaard, 1988). However, the Cooma Granodiorite differs from the “batholithic” cordierite granites in having systematically lower Na2O, CaO, and Sr, but higher SiO2 and ASI (Figure 16). It also shows different trends for some elements, particularly for TiO2, K2O, and Zr, which mirror the compositional variation of the Ordovician turbidites rather than the array defined by most cordierite granites. The Cooma Granodiorite also lacks the microgranular enclave population present in the lower silica cordierite granites and has the highest strontium and lowest neodymium isotopic ratios (Figure 14). Some other model seems to be required for most of the LFB cordierite granites. There are two potential resolutions to this problem. One is to suggest that cordierite granites were derived from more feldspathic (i.e., less weathered) sedimentary rocks with higher Na2O and CaO contents than most of the quartz-rich Ordovician turbidites (Chappell et al., 2000). Additional support for this is that the unfractionated cordierite granites plotted on Figure 11 only show compositional overlap with the least weathered Ordovician sediments that have the lowest Rb/Sr ratios, but highest Sr/Nd and Sr/Eup (and Eu/Sr, see Figure 20(a)). Such relatively fertile protoliths have yet to be recognized in the LFB and thus would need to be entirely concealed beneath the ubiquitous Ordovician turbidite sequence. The large population of inherited zircons with Ordovician to Late Cambrian ages in the cordierite granites suggest a dominantly Early Paleozoic, rather than Proterozoic age for such a protolith (e.g., Keay et al., 1999). Some evidence for the existence of a more fertile source layer at depth is the generally higher CaO and Na2O contents of metasedimentary enclaves in the cordierite granites compared to the exposed turbidites (Figure 16) (White et al., 1977; Maas et al., 1997; Chappell et al., 2000). Yet the possibility that some of these are restitic (calcic plagioclaserich) or modified in some way by the host magma cannot be excluded. The paucity of metasedimentary enclaves with between 60 – 75% SiO2 is striking, as this is consistent with the most fertile parts of the sequence having melted to form the granitic magma and thus being unavailable for sampling as enclaves. It is possible that some of the peraluminous “microgranular” enclaves that fall within the 60 – 75% silica range represent such source material. If the batholithic cordierite granites formed by mobilization of a fertile metasedimentary source

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 395 in analogous fashion to the Cooma Granodiorite (White et al., 1974), then the protolith must average at least 3% CaO, 3.5% K2O, and 150 ppm Sr at 66 –67% SiO2, but contain less than 2% Na2O. Paradoxically, sedimentary rocks of this specific composition do not outcrop on the Australian mainland (Collins, 1998). In fact, in the context of global sedimentary compositions, such a rock seems abnormally calcic and sodium-poor; Phanerozoic graywackes of any tectonic setting with .3% CaO have Na2O/K2O ratios exceeding unity, a reflection of their detrital feldspar component (see Taylor and McLennan, 1985 and the compilation of Condie, 1993) However, it is possible that the bulk metasedimentary package undergoing melting was of this composition (i.e., represents carbonates mixed with graywackes). In the absence of a suitable exposed metasedimentary precursor, the second approach is to invoke mixing with an additional magmatic component (e.g., Gray, 1984, 1990; Collins, 1996; Keay et al., 1997). This has the advantage of explaining the pronounced isotopic variation within cordierite granites, since restite separation and fractional crystallization are isotopically closed system processes. Collins (1996) argued that the lower silica cordierite granites formed in the mid-crust by large-scale hybridization between partially melted Ordovician sediments, approximated by Cooma-type magmas (,72% SiO2), and hotter, ascending tonalite magmas (,60% SiO2), the proportions being 60– 70% Cooma Granodiorite and 30 – 40% tonalite. However, hornblende tonalites are volumetrically minor in the LFB, and the specific low titanium, low strontium composition required to satisfy the chemical modeling is anomalous amongst hornblende granites. Furthermore, microgranular enclaves in the cordierite granites define trends that converge towards genuinely basaltic compositions, implying the involvement of mantle-derived, rather than tonalitic, magmas (Figure 16). Yet the geochemistry of cordierite granites also seems to prohibit a significant material contribution from basaltic magmas. It is unlikely that the clustered, lowest-silica cordierite granites were generated by simple mixing between basalt and Cooma type diatexite, since the low Na2O of the granites requires the basalts to have unrealistically high Na2O contents (,4% at 50% SiO2). The trends of cordierite granites cannot represent mixing lines between felsic crustal melt and mafic magma, since the cordierite granites become increasingly peraluminous as they become more mafic (Figure 16), which requires a strongly corundum-normative basalt (White and Chappell, 1988). The decrease in initial 87Sr/86Sr of cordierite granites as silica increases (Figure 18(b)) is also

inconsistent with such bulk mixing. This discussion highlights that although open-system processes are implicated in the genesis and compositional evolution of cordierite granites, these were evidently more complex than binary crust-mantle mixing. Hornblende granites. Field and chemical evidence, especially the projection of geochemical trends towards coeval gabbros, confers a pivotal role for mafic magmas in the production of the LFB hornblende granite suites. However, many of these plutons have “crust-like” traceelement patterns similar to the cordierite granites (Figure 17), and they contain inherited zircons that have age populations identical to that of the Ordovician turbidites (Williams et al., 1992; Williams, 1995). Thus the hornblende granites also incorporated a sedimentary component (Gray, 1984, 1990; Collins, 1996, 1998; Keay et al., 1997). The inherited zircons are particularly significant, since, assuming that their survival does not reflect sluggish dissolution kinetics, they also suggest that the hornblende granites were not liquid compositions (Chappell et al., 2000), and they had magmatic temperatures lower than that required for the resorption of entrained zircon crystals (i.e., below , 800 8C). In view of this, the hornblende granites cannot have evolved simply by fractional crystallization from basaltic magma, and this is also ruled out by their evolved isotopic compositions relative to mantle values. Instead, as with cordierite granites, it seems that the formation of hornblende granites was linked to crustal melting.

3.11.4.4.5

A petrogenetic model

The spatial and temporal association, as well as the chemical similarities, suggests that generation of the cordierite and hornblende granites in the LFB was a coupled process and involved similar source materials. The preferred model entails the successive emplacement of thick basaltic sills in the lower crust at the onset of Lachlan orogenesis, these being overlain by a thick collage of Ordovician sediments and their oceanic crust substrate. The existence of large amounts of mafic material in the lower parts of the LFB crust is indicated by lower crustal xenolith studies (Chen et al., 1998) and seismic profiling (see Anderson et al., 1998), and confirms that massive basaltic intrusion, and thus crustal growth, accompanied granite genesis. Initially, the sills solidified rapidly, but as the process continued, and the lower crustal region was progressively heated, the larger bodies would have remained partially molten for longer periods, to undergo protracted

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Granitic Perspectives

crystallization. The heat (and perhaps fluid transfer) from these crystallising mafic magmas would have induced fusion of the overlying metasedimentary rocks, and any intercalated screens of such material. The most fertile parts of the sedimentary sequence would yield the greatest proportion of initial partial melt, and segregation of this was possibly responsible for the formation of the felsic cordierite granites. These “minimum melts” would possess relatively low 87Sr/86Sr, given the feldspar-rich nature of their protoliths. As the lower crustal section continued to heat up, melting would proceed to less fertile lithologies, and this would include fragmentation and assimilation of the more refractory, calcareous horizons. The trends to lower silica of the cordierite granites are thus interpreted in terms of increasing degrees of partial melting and possibly restite contamination, the trajectory of which is a vector towards the bulk composition of the turbidite sequence. The increasing restite cargo is strongly suggested by the greater proportion of inherited zircon (Williams, 1995) and metasedimentary enclaves in lower silica cordierite granites. Samples of the clustered compositions were therefore not liquid compositions but probably contained suspended crystals, both refractory and the byproducts of partial melting reactions. In support of this, liquids resembling the cordierite-rich granite compositions have not yet been reproduced experimentally from metasedimentary protoliths under any P – T-aH2O conditions (Montel and Vielzeof, 1997; Patin˜o Douce, 1999). The increase in 87Sr/86Sr in granites having SiO2 contents of ,74% to 68% is consistent with greater incorporation of metapelitic rocks, the radiogenic signature of which (87Sr/86Sr , 0.726 at the time of granitic generation, McCulloch and Woodhead (1993)) is likely to swamp low 87Sr/86Sr contributions from the calc-silicates. However, the very high CaO of the latter (11 – 13%) largely controls the accompanying trend to higher CaO. The numerous metasedimentary enclaves and micaceous clots in the lower silica cordierite granites comprise relatively refractory material and may have been entrained from the metasedimentary source regions (Chappell et al., 1987). Alternatively, they could have been derived from interaction with Cooma-type diatexite during passage through the middle crust, as the scattering of some cordierite granite compositions to high ASI values suggests that this was important locally (see Figure 16). Periodically, residual mafic liquids from the underlying sills would percolate upwards to mix with the overlying partially melted sediments and crustal magmas. This is considered responsible for driving some cordierite granites towards slightly

more mafic, less peraluminous compositions, and for the observed decrease in initial 87Sr/86Sr below 68% SiO2. Most microgranular enclaves were conceivably also dispersed by this process, and their complex chemical variation can be interpreted as reflecting hybridization between variably differentiated basaltic liquids and cordierite granite magma. Given the trends to low Na2O and high ASI, some enclaves may also have formed by hybridization between basaltic liquids and infertile Ordovician sediments, to be subsequently fragmented and entrained by the ascending granitic magma. Continuation of the underplating and crystallization process would eventually remelt the earlier, solidified and crustally contaminated basaltic sills, or their evolved, differentiates, and this is considered responsible for the formation of the hornblende granites. The extraction and emplacement of the initial, minimum melts resulted in the felsic hornblende granite plutons that dominate the eastern LFB. More mafic compositions were produced by increasing degrees of melting and also by replenishment of the partially melted sills by more primitive basaltic liquids. The latter process is recorded by the mafic enclaves that are abundant in the lower silica samples. The composition of each derivative hornblende granite suite is therefore a function of the geochemistry of the participating basaltic magma and the amount of crustal material (melt or bulk sediment) assimilated by the crystallizing sills and later remelts. For example, granitic rocks of the Moruya suite in the easternmost LFB have a large compositional range (60 – 74% SiO2) but are uniformly isotopically primitive (1Nd , þ4) showing that contamination by sedimentary materials was minor (,6%, assuming the isotopic ratios for the basalt and sediment end-members indicated in Figure 14). The isotopic systematics and trace-element patterns of plutons from the eastern edge of the Bega batholith also seem to preclude a substantial sedimentary component (see Figure 17). This has been attributed to the paucity of sedimentary rocks at the depth of granite generation in the eastern LFB, possibly reflecting thinning of the turbidite blanket outboard of the continental margin (Collins, 1998). The distinctively high Na 2O and Sr that characterizes hornblende granite suites of the eastern LFB was clearly imparted by the ancestral mafic magmas, as these features are evident within the coeval basaltic dikes analyses presented by Griffin et al. (1978). In this respect, as first pointed out by Chappell and Stephens (1988) and subsequently emphasized by Keay et al. (1997) and Collins (1998), the eastern Lachlan granites resemble the sodic plutons of

Generation of High Silica Continental Rocks: Proterozoic to Phanerozoic Granitic Associations 397 convergent continental margins, which have largely basaltic protoliths and minimum input from evolved supracrustal materials (see discussion in Section 3.11.3.1.1). In contrast, hornblende granites of the Berridale Batholith further west are more potassic and isotopically evolved (ranging to 1Nd , 28.1, initial 87Sr/86Sr ,0.712; McCulloch and Chappell, 1982). These rocks are associated with cordierite granites, indicating that a substantial thickness of Ordovician sediments existed at mid-crustal depths, some of which were incorporated by the hornblende granites, or their mafic precursors. This accounts for the population of inherited zircons and similar trace-element patterns to the cordierite granites. Unlike other mixing models, as the mafic and felsic components are genetically related, this scenario has the unique advantage of being able to explain the chemical similarities shown by hornblende granites at opposite extremes of the compositional range. This observation has been previously used as a major objection to the viability of magma mixing in the LFB (e.g., Chappell, 1996a). The compositional trends of hornblende and cordierite granites are thus explicable in terms of magma mixing and partial melting, respectively, with an ancillary role for restite entrainment in cordierite granites. Fractional crystallization, as advocated for cordierite granites by Collins (1996), is less favored since these rocks were clearly not liquids at the mafic end of their compositional spectrum. Fractionation is also not compatible with the reasonably flat trends for strontium, barium, and rubidium, at least to 74% SiO2. The role of fractionation in the formation of the hornblende granites in general is difficult to assess, not least because fractionation trends will parallel partial melting trends, and awaits more isotopic data. However, the sharp increases in Rb/Sr and Rb/Ba, and pronounced downward inflections in strontium, CaO, Zr, La, and Ce at greater than 74% SiO2 (Figure 15(a)) suggest that the final stage in the evolution of both LFB granite types did involve fractional crystallization, controlled predominantly by feldspars and accessory minerals. For cordierite granites, this is consistent with the sharp increase in ASI in the most felsic rocks (i.e., .74% SiO2), since this is elevated by the removal of feldspars (ASI ¼ 1.0) from peraluminous magmas. The steep P2O5 increase is coupled to this increase in ASI, since apatite solubility increases drastically with peraluminosity (Pichavant et al., 1992). In contrast, the decrease in P2O5 shown by hornblende granites indicates that apatite saturation was reached for these less peraluminous magmas (Chappell, 1999). An implication of this is that the most evolved compositions in granitic suites should not be used as end-members in mixing

models, as any fractionation will produce spurious results, especially on incompatible/compatible element ratio plots. In summary, the generation and evolution of the contrasting granitic types in the LFB appears to have been a three stage process that was driven by incursion of mantle-derived magmas into the crust. For the hornblende-bearing granites, the stages involved (i) remelting and/or prolonged fractionation of crustally contaminated basaltic sills to produce silicic minimum melts; (ii) increasing degrees of partial melting, combined with mixing with fresh basalt to form more mafic compositions; and (iii) fractional crystallization. For the cordierite granites, these stages are: (i) varying degrees of anatexis of Ordovician sediments; (ii) hybridization between these crustal magmas and basalt to produce the lowest silica compositions; and (iii) fractional crystallization at the highest silica contents. Greater insight into the way that the mantle- and crustally derived source components were combined during generation of the cordierite and hornblende granites awaits more systematic documentation of within-suite isotopic variation. Studies that attempt to unravel the magmatic evolution of individual plutons, for example by tracking changes in hafnium isotopic composition across magmatically zoned zircon grains (e.g., Griffin et al., 2002), should prove to be particularly valuable. 3.11.4.4.6

Relevance for crustal evolution

If the model outlined above is valid, silicic magmatism in the LFB involved net crustal growth, as juvenile mantle-derived liquids, or their differentiates were instrumental in the formation and compositional evolution of hornblende granites, and, to a lesser extent, the cordierite granites. The amount of new crust generated is estimated by determining the overall mantle component present within both granitic types, and this is best done isotopically, since the trace-element ratios are poorly constrained for the potential basaltic end-members. For the purposes of modeling it is convenient to assign a common depleted mantle-like endmember to all hornblende and cordierite granites (McCulloch and Chappell, 1982). This therefore represents the minimum case for crustal growth, since the mantle-derived end-member could be significantly more evolved than this, especially if generated in a subduction-related setting. Inspection of Figure 14 reveals that the average hornblende granite isotopic composition (initial 87Sr/86Sr ¼ 0.707, 1Nd 23) falls on a mixing curve between its inferred depleted mantle (basaltic) and crustal end-members, and corresponds to ca. 70% of the mantle component. For cordierite granites, the plotted composition

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represents the average low silica granite (those with ,68% SiO2), for which an initial 87Sr/86Sr of 0.7135 and 1Nd 28 is inferred; such a composition is modeled as ca. 10% depleted mantle component. Assuming that hornblende and cordierite granites occur in equal proportion and constitute 30% of the LFB, and given that the low silica samples thought to incorporate a mantle component comprise ,20% of all cordierite granites, this indicates that ,10% of the presently exposed LFB crust was newly generated during orogenesis, as manifested within granitic intrusions. This estimate is admittedly crude, and serves for illustrative purposes only, but should become more refined as future studies pinpoint the nature of the mantle-derived component and establish the range of isotopic variation shown by the various granitic suites. The percentage of new crust will increase if it can be shown that more enriched basaltic rocks were involved in granite petrogenesis. Although hornblende granites contain a greater mantle component than the cordierite granites, the chemical signature of this ingredient is “swamped” by that of the Ordovician turbidites. This explains the similar trace-element ratios of some hornblende and cordierite granites, despite the inferred differences in petrogenesis. Differentiation of the newly formed LFB crust was achieved predominantly by intracrustal melting, involving both fusion of the underplated basaltic sills, and mixing between basaltic magmas and relatively evolved magma derived from the pre-existing supracrustal rocks, with the superimposed effects of fractional crystallization. This reinforces the point that crustal growth and intracrustal melting are likely to be coupled during orogenesis, the latter being an inevitable consequence of the heat and (in a subduction setting) volatile transfer from the emplacement of mantle-derived magmas. However, although similarly related to crustal anatexis, the style of crustal differentiation associated with granitic magmatism in the LFB differs from that of strongly peraluminous leucogranites in three main respects. First, approximately half of the plutons (hornblende granites) are derived from juvenile mantle-derived protoliths, rather than pre-existing sedimentary rocks. Second, in respect of the LFB cordierite granites, the crustal melts also incorporate basalt and some restitic material, partly explaining their more mafic compositions. Hence, they have higher REE contents and have trace-element patterns that more closely resemble the upper crust (Chappell, 1996b). Yet as with the hornblende granites, strontium, and phosphorus contents are lower than the bulk and upper continental crust compositions. Thirdly, heat producing elements are more efficiently transferred to the upper crust via cordierite granite magmas than in

leucogranites. This does not reflect greater solubility of refractory zircon and monazite, but rather the physical entrainment of such minerals, probably enclosed within major phases, by the magma (Chappell, 1996b). Sandiford et al. (2002) emphasize that scavenging of heat producing elements from the lower parts of the crust will induce long term cooling of these regions, substantially increasing the strength, and also density of the lithosphere. The density increase is predicted to trigger isostatic subsidence, and such subsidence is manifested as thick sedimentary basins in northern Australia that formed following the emplacement of plutons enriched in heat producing elements (Sandiford et al., 2002). The development of Late Devonian sedimentary basins in the LFB immediately after plutonism implies that thermal subsidence related to the redistribution of K –Th – U may also have been important in this area.

3.11.5 CRUSTAL DIFFERENTIATION THROUGH EROSION AND SEDIMENTATION 3.11.5.1 The Link Between Rb/Sr Ratios and Granitic Sources The discussion this far has highlighted that granite magmatism is a key process in the differentiation and evolution of the continental crust. However, there is no simple match between the composition of common granites and that of the average upper crust, and so some other process(es) must be involved. This section compares the effects of granite formation, erosion and sediment formation on the evolution of the upper crust as expressed through its Rb/Sr ratio. This ratios is apposite because it is the parent/daughter ratio for the Rb – Sr decay scheme, and so it is constrained by measured 87Sr/86Sr values, and it is sensitive to the effects of residual plagioclase (which is only really stable within the crust) and weathering and erosion. Detailed studies of granitic rocks in a number of orogenic belts have identified broad positive arrays between the Rb/Sr ratios of individual granite samples and the time-integrated Rb/Sr ratios of their source regions inferred from their model neodymium ages. The model neodymium ages of the granites are used to estimate the average age when their crustal sources were extracted from the mantle, and then the strontium isotope ratio of the mantle at that time, and the initial strontium isotope ratios of the granites are used to calculate the time-integrated Rb/Sr ratios of the granites’ source regions. These are then compared with the measured Rb/Sr ratios of

Crustal Differentiation Through Erosion and Sedimentation

Rb/Sr (rock)

10

1 Idaho Bath. Hbl granite Crd granite GRC LG Himalayan Damara

0.1

0.01 0.01

0.1 Rb/Sr (source)

1

Figure 19 Plot of Rb/Sr in granitic rocks versus that calculated for their putative source regions to illustrate the degree of crust differentiation using Rb/Sr ratios (the 1:1 line is dotted). See text for full explanation. The data are from: Idaho Batholith, Clarke (1990); LFB hornblende (Hbl) and cordierite (Crd) granites, as for Figure 18; Glenelg River Complex leucogranites (GRC LG), Kemp (2001); Himalayan leucogranites, Inger and Harris (1993); Pan-African granitic rocks of the Damara Orogen, Namibia (Hawkesworth and Marlow, 1983; McDermott et al., 1989; McDermott and Hawkesworth, 1990).

individual granite samples in Figure 19. This illustrates that more evolved, and hence high Rb/Sr granites tend to be derived from source regions that are also more evolved, again as reflected in their higher source Rb/Sr ratios. On the basis that the longer rocks remain in the crust, the more likely they are to be involved in further magmatic or sedimentary events (e.g., Davies et al. (1985) concluded that the upper crustal rocks of the British Isles were reworked on average every 600 Ma), it might be predicted that the timeintegrated source Rb/Sr ratios would increase with their model neodymium ages. However, in practice there is no simple link between the measured or the source Rb/Sr ratios and the age of the source regions for the granites, as reflected in their model neodymium ages. This suggests that the source Rb/Sr ratios primarily reflect the rock types in the granite source regions, rather than the length of time the source rocks, or their precursors, have been in the crust. 3.11.5.2 Resolving the Effects of Igneous and Sedimentary Processes Weathering and erosion at the Earth’s surface preferentially removes those elements that are water soluble (e.g., strontium, calcium, barium, magnesium) from the continental mass, and enriches elements that are adsorbed onto clay minerals (e.g., rubidium, REE3þ). One way to isolate the effects of igneous and sedimentary processes is therefore to use trace-element ratios

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involving europium and strontium. These elements are both sited in plagioclase but they behave differently in igneous and sedimentary environments. During fractional crystallization (or partial melting) within the crust, strontium and europium are both partitioned into plagioclase, and so their concentrations are lower in the coexisting liquids. Europium is also depleted relative to the trivalent REE, leading to the development of progressively larger negative europium anomalies (i.e., Eu/Sm, which is a proxy for europium anomalies, will decrease) as differentiation proceeds. However, the degree to which europium is partitioned into plagioclase depends on the proportion of Eu2þ and Eu3þ, and therefore on the oxygen fugacity (Drake and Weill, 1975). In contrast, strontium and europium are decoupled during sedimentary processes. This is because the soluble strontium is released to the oceans during weathering of feldspathic rocks, but europium is retained in the residual weathering profile, as Eu2þ is oxidized to Eu3þ and thus behaves like the other trivalent, immobile REE. Weathering therefore cannot produce the depletion in europium relative to samarium that fingerprints magmatic differentiation, and this can be used to assess the extent to which strontium depletion in upper crustal rocks is due to magmatic or sedimentary processes. To explore these effects, Figure 20 shows plots of (a) Rb/Sr versus Eu/Sr and (b) Eu/Sm versus Eu/Sr for various types of granitic and sedimentary rocks, compared to estimates of crustal compositions from Table 1. It is unclear how representative these rocks are of upper crustal processes, but they provide a useful framework for discussion, particularly since the igneous rocks include relatively “primitive” TTG-like plutons (i.e., Idaho Batholith) and highly evolved A-types, and thus encompass the geochemical spectrum of granitic compositions in the crust. On Figure 20(a), the igneous and sedimentary rocks plot in an overlapping positive array, since the Rb/Sr ratio is elevated by weathering (rubidium is retained in clays, and strontium is lost to sea-water) and igneous fractionation (rubidium is incompatible and strontium is compatible in plagioclase). In sedimentary rocks, the increase of Eu/Sr with differentiation reflects the removal of strontium to the oceans, whereas for igneous rocks it indicates that strontium is more compatible than europium in the residual or crystallizing plagioclase feldspar. This is highlighted in Figure 20(b) since europium and samarium are not fractionated during weathering, and so the igneous and sedimentary trends tend to diverge. Two other points are worth noting. First, apart from the A-type granites, the rocks with elevated Rb/Sr and Eu/Sr ratios in Figure 20 tend to be sediments and sediment-derived granitic rocks.

400

Granitic Perspectives 1000

100 Idaho Bath. Andes

Rb/Sr

10

LFB Hbl granite LFB Crd granite

1

BC

Damara grans. A-types Damara seds.

LC

Cambrian seds. LFB sediments

UC 0.1

0.01 0.001 (a)

0.01

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1

Eu/Sr

1 LC Sedimentary processes

Eu/Sm

BC

0.1

Igneous processes 0.01 0.001 (b)

0.01

0.1

1

Eu/Sr

Figure 20 Plots of (a) Rb/Sr versus Eu/Sr and (b) Eu/Sm versus Eu/Sr for the reference igneous and sedimentary suites referred to in this chapter, compared to estimates of the continental crust (UC, upper crust; BC, bulk crust; LC, lower crust, all connected by tie lines). The contrasting trends produced by igneous and sedimentary differentiation processes in (b) are arrowed. Data sources are as for previous figures, with the LFB granites and Cambrian Delamerian Orogen sediments (southeastern Australia) from Kemp (2003), unpublished data), and the granitic and sedimentary rocks of the Damara Orogen from McDermott and Hawkesworth (1990) and McDermott et al. (1996, 2000).

This further emphasizes the importance of sedimentary processes in the development of such fractionated Rb/Sr and Eu/Sr ratios. Second, in the context of the LFB, the Ordovician turbidites have progressively lower Eu/Sr ratios than the feldspathic Cambrian sediments from which they were partly derived. Thus, it also appears that the displacement to higher Eu/Sr is linked to the number of weathering and erosion cycles involved in the formation of particular sedimentary rock suites. In the examples in Figure 20, the reduction in strontium content for the most evolved Ordovician sediments is nearly 10-fold, and this clearly results in an associated increase in Rb/Sr, and hence with time in strontium isotope ratios.

3.11.5.3 Relevance for Crustal Differentiation The fractionated igneous and sedimentary rocks in Figure 20 have Rb/Sr and Eu/Sr ratios that are much higher than those in current estimates for the upper continental crust. The latter can be constrained by the strontium isotope ratios of continental run-off (,0.712), and its model neodymium age (,1.8 Ga). According to this method, a minimum timeintegrated upper crustal Rb/Sr ratio of 0.14 is indicated. Another striking feature of the data in Figure 20 is that the upper, lower and bulk continental crust compositions all have similar Eu/Sr ratios and thus define a distinct, near-vertical trend that is

Synthesis and Avenues for Future Study separate from the igneous and sedimentary arrays. There are several potential interpretations for this. It might merely reflect the methods employed to estimate the upper and lower crustal averages. The upper crust represents a mixture between sediments and intermediate to felsic igneous rocks (note how close the upper crustal composition plots to the reference suite array in Figure 20), whereas the lower crust inevitably combines xenolith data from both intraplate and destructive margin settings, not necessarily in representative proportions; the bulk crustal composition is constrained to lie between these extremes. The dispersion in crustal compositions on Figure 20 could therefore be synthetic and petrogenetically meaningless. Alternatively, if the chemical variation between the crustal components results from the differentiation or “unmixing” of a bulk starting composition, the data suggests that neither igneous fractionation nor weathering processes can be wholly responsible for such differentiation, since strongly increasing Eu/Sr is a signature of those processes. In any case, it is clear from Figure 20 that the upper crustal composition has sufficiently low Rb/Sr and Eu/Sr to preclude a significant contribution from the continental sediments, in contrast to the implications drawn from the granite and upper crust trace-element patterns in Figure 5. One resolution to this conundrum could be related to a shift in the oxidation state of europium (i.e., the proportion of Eu2þ to Eu3þ) either through time, or in different tectonic settings (Carmichael, 1991). The marked increase in Eu/Sr in igneous suites indicates that DSr was much larger than DEu in the fractionating plagioclase. However, DEu in plagioclase is sensitive to oxygen fugacity ( fO2), and it is very low in oxidizing conditions where europium exists as Eu3þ (Drake, 1974). The igneous rocks plotted in Figure 20 can be inferred to have formed under reasonably oxidizing conditions, since many of them originate above subduction zones, where the mantle has been modified by the introduction of hydrous fluids from the subducted slab, or they derive from, or have interacted with, recycled sedimentary rocks in the deep crust. However, in reducing conditions, DEu increases until, at fO2 ,10 – 12.5 bars, it approaches values similar to that of DSr. Plagioclase formed under these conditions will therefore not fractionate europium from strontium, and its removal results in vertical arrays on Figure 20, as is shown by the continental crust. This raises the intriguing possibility that the differentiation of the continental crust was primarily achieved under relatively reducing conditions, such as existed in the Archean period where a CO2-rich atmosphere prevailed, or in intraplate settings. The latter would be marked by magmas with distinctive

401

trace-elements patterns, and in particular no negative Nb –Ta anomalies. Such magmas did contribute to the generation of new crust, they cannot be the dominant component, and so the apparent lack of Eu/Sr fractionation in the crustal compositions may largely reflect processes in the Archean. Independent evidence for a reducing environment at this time includes the presence of banded iron formations, uranium placer deposits, high Th/U ratios in igneous rocks, non-mass dependent sulfur isotope fractionations (Farquhar et al., 2000; see also Chapter 4.04) and the lead isotope composition of the mantle (see Elliott et al., 1999). Under such reducing conditions, intracrustal melting can generate the observed differentiation of the continents, the important point being that the residual plagioclase contained substantially more europium than that of the magmatic reference array in Figure 20. Recycling of small amounts of the residues of melting, approximated by the estimated lower crustal composition, can explain the displacement of the bulk crust from the magmatic array on Figure 20. The fundamental implication from this reasoning is that melting and weathering processes operating in the outer part of the post-Archean Earth have contributed relatively little to the bulk differentiation of the continental crust, consistent with its average age of ,1.8 Ga.

3.11.6 SYNTHESIS AND AVENUES FOR FUTURE STUDY After decades of effort, we appear to have arrived at robust estimates for the composition of the continental crust that are corroborated by diverse geochemical, isotopic and geophysical approaches. The greatest source of uncertainty is that models for the bulk crust depend on estimates for the lower crust, which are less tightly constrained. However, it is well recognized that although the continental crust has grown throughout Earth’s history by material flux from the underlying mantle, the bulk composition is not that of any common primary mantle-derived magma, either basaltic or andesitic. The discrepancy is greatest with elements that behave compatibly during deep (magnesium) and midto shallow-level crustal processes (strontium, europium) and thus all current models require preferential recycling of these elements from the continental crust back into the mantle. Isotopic studies have established that the continental crust is old, in that more than half of it was formed by the end of the Archean. The distinctive igneous components of Archean cratons, coupled with evidence that the crust at this time was overall less mafic, provide clues that crustal growth mechanisms during this period

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Granitic Perspectives

differed significantly from those operating today. The earliest continental nuclei, the Archean TTG, represent high-pressure, second-stage melts of juvenile mantle-derived basalts, analogues of which are rare in post-Archean orogens. However, the tectonic setting in which the TTG, and indeed new crust, were generated in the Archean remains uncertain. This inhibits our understanding of the way in which the crustal protoliths were extracted from mantle sources in the Archean, and how selected components were returned to the mantle. The apparent periodicity in crustal growth, the bimodal nature of Archean magmatism, and the lack of a consistent subduction signal in certain greenstone belt volcanic rocks, are strong evidence for the role of plume-related magmatism. Alternatively, compressive tectonic regimes that may well reflect subduction processes are indicated by the depth of melting of the TTG, and the time interval of 50 – 80 Ma between the generation of mafic and felsic magmas. Subduction also offers a convenient mechanism for dislocating TTG magmas from their voluminous ultramafic residues, which are recycled into the mantle via foundering slabs of oceanic lithosphere. A satisfactory compromise might therefore involve models whereby the basaltic protoliths to the TTG are formed in a plume-related environment, and are subsequently remelted to form the silicic magmas during subduction of this young mafic crust. Elemental ratios involving the HFSE can in principle be linked to the behavior of key mineral phases and so place some constraints on the nature of the crust-forming process through time. However, the evidence from HFSE systematics may sometimes be ambiguous. For example, negative Nb – Ta anomalies are widely used to identify subduction related magmatism in ancient orogenic belts, and to distinguish it from what is often inferred to be plume-related intraplate melt generation. However, negative Nb –Ta anomalies may also be generated during intracrustal melting and magma differentiation, and so mask evidence for the tectonic setting of crust generation. In this chapter, we have emphasized that felsic magmatic rocks, the buoyant building blocks of the continents, show dramatic secular variation, and that this can be related to changing thermal and atmospheric conditions in the Earth. Regardless of the tectonic setting, the thermal conditions necessary for remelting basaltic rocks in the garnet stability field were rarely attained after the Archean. Furthermore, key chemical indices, especially the development of a europium anomaly in the upper crust, indicate that mid to shallow level intracrustal differentiation under high fO2 became more common in the postArchean Earth. The estimated crustal structure suggests that elements with similar characteristics, such as niobium, tantalum, and titanium, and

strontium and europium, were fractionated during formation of the upper crust, such that it is relatively enriched in Nb –Ta and Sr, but depleted in titanium and europium compared to the bulk composition. Crustal differentiation leading to an evolved upper crust is conventionally attributed to partial melting and granite formation, augmented by erosion and weathering, with the implicit assumption of “closed system” behavior. Furthermore, one might predict that the degree of crustal differentiation should increase progressively with time, given the abundance of broadly granitic rocks emplaced into the upper crust of all time periods. However, from an examination of the chemical fractionation associated with granitic magmatism and sedimentary processes, the following points emerge. (i) In the absence of heat transfer from mantlederived magmas, crustal melting produces silicarich and strongly peraluminous leucogranites by low temperature, muscovite melting reactions. The extreme chemical fractionation associated with this “haplogranitic” style of magmatism cannot drive crustal differentiation, though it is potentially important in enriching the upper crust in the most incompatible elements (e.g., rubidium and potassium). Regardless of the presence or absence of fluid during melting, the low REE and HFSE of these rocks is buffered by the low solubility of refractory accessory minerals, chiefly zircon and monazite. The heat producing elements uranium and thorium may not therefore be mobilized in this style of granite magmatism. (ii) From the Late Archean onwards, the formation of relatively mafic granites (SiO2 , 70%) with elemental ratios more like the continental crust required higher temperature melting reactions and direct heat input from mantle-derived magmas. Chemical and isotopic considerations from specific case studies demand that juvenile mantle-derived magmas have also made a material contribution to granites generated from a range of crustal protoliths; i.e., the Moho has acted as an open interface during post-Archean granitic generation. Granitic rocks should therefore be properly viewed as the products of multistage processes, often involving materials from more than one source and from different geochemical reservoirs. It therefore seems sensible to retain classification schemes for these rocks that are descriptive, rather than generic. (iii) At the same silica content (,66% SiO2), the granites of oceanic arcs, continental arcs, and large, probably subduction-related granitic provinces that also include postcollisional magmas (LFB), do not match the estimated upper crustal composition. In arc granites, Nb –Ta and Ti are coupled, and so these magmas cannot enrich the upper crust in Nb – Ta relative to Ti, unless Nb –Ta

Synthesis and Avenues for Future Study contents are elevated by an additional process (see point 5 below). Arc granites do however, have high Sr/Nd, and they therefore contribute to maintaining the moderately high strontium contents in the upper crust. Granites of the LFB have high rubidium, potassium, thorium, HFSE, and REE contents, reflecting variable contributions from mature sediment. Consequently, however, these granites alone do not have sufficient strontium to account for its abundance in the upper crust. (iv) A-type granites appear mostly to have formed by relatively shallow level melting or fractionation of basalt in the absence of garnet, and thus they represent the other end of the compositional spectrum to the TTG on some chemical diagrams (e.g., Figure 20). These magmas represent the clearest magmatic mechanism for enriching the upper crust in the HFSE relative to titanium, but are associated with extreme strontium, barium, phosphorus, and europium depletion. To understand the role of these granites further in crustal evolution, we need to ascertain why there was a major pulse in A-type generation in the Proterozoic. (v) Following from points (iii) and (iv) above, the best fit to the average composition of the upper crust is a , 75 : 35 mixture of intermediate to silicic arc magmas (which have relatively high strontium, but low titanium and HFSE) and sedimentary rocks (low strontium and titanium, but high HFSE) (Figure 21). This is consistent with the positioning of the upper crustal composition within the igneous and sedimentary array on Figure 20, and tends to diminish the importance of nonarc magmas for determining the upper crustal composition. The intermediate subduction-related magmas in this model may have evolved from basaltic liquids by high-pressure fractionation and contamination at the base of thickened crust, as suggested for some Andean andesites by Davidson and Arculus (2001). (vi) Granites with high Rb/Sr ratios are usually attributed to partial melting processes involving mica breakdown and/or later plagioclase fractionation, the latter being most clearly the case for the A-types. However, we contend that the Rb/Sr ratios of granites also strongly reflect the Rb/Sr ratios, and hence the degree of evolution, of their source rocks (Figure 19). This argument is underpinned by isotopic evidence and thus makes no assumption about the melting reactions involved in granite genesis. (vii) Igneous differentiation in the crust results in decreasing Eu/Sm (proxy for Eu/Eup) and its effects may therefore be distinguished from those of weathering and erosion, where this ratio is unchanged. Both sedimentary and igneous processes are associated with markedly increasing Eu/Sr, reflecting strontium loss to the oceans in the former, and DSr q DEu in the latter.

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10

Granite-sediment mixing

1 average PRB av. GRC schist upper crust PRB-schist mix 0.1

Bulk crust normalized

Rb Ba Th U K Nb La Ce Pb Sr Nd P Zr Hf Ti Y

Figure 21 Bulk crust normalized diagram showing simple mixing between the average granitic rock from the subduction-related Peninsula Ranges Batholith, California (PRB, from Silver and Chappell, 1988, with 64.6% SiO2), and the average of 92 metasedimentary rocks from the Cambro-Ordovician Glenelg River Complex (GRC, from Kemp, unpublished data, with 71.1% SiO2) in the proportion 75% granite, 35% sediment. The resultant composition has 66.9% SiO2, comparing favorably with estimates of the upper continental crust (66.6% SiO2), and closely mimics the upper crust for most trace elements. Phanerozoic arc magmas are most suitable for this type of modeling, owing to their high Sr contents, a reflection of the preferential transport of this element by water-rich fluids from the subducted slab to mantle wedge during magma generation (Pearce and Peate, 1995). Less satisfactory fits are obtained for the TTG, since the resultant mixtures have much lower HREE and Y than the contemporary upper crust.

Strikingly, however, estimates of the upper, middle and lower continental crust do not follow such differentiation trends, and in particular there is relatively little fractionation of Eu/Sr with increasing Rb/Sr. If this is not an artifact of how the crustal estimates were derived, it implies that differentiation of the bulk crust took place under different conditions than those prevailing during the formation of the reference magmatic suites in Figure 20. We tentatively suggest that crustal differentiation may have been dominated by conditions in the Archean when oxygen fugacity was less, and thus the relative partitioning of europium and strontium into plagioclase was different than in post-Archean igneous processes. In conclusion, a number of arguments presented in this chapter have highlighted the pivotal role of Archean processes in shaping the compositional structure of Earth’s continental crust. In view of this, and despite the spectacular mode of crustal growth witnessed in modern convergent margin and oceanic island volcanoes, it may be that for crustal studies we need to increasingly focus our attention on the distant past, as the key to understanding the present.

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ACKNOWLEDGMENTS The ideas presented in this manuscript have been focused and strengthened by the advice and thorough review by Roberta Rudnick, and we also appreciate her assistance with editorial matters. Helpful discussions with Simon Turner, Jon Blundy, and Bruce Chappell, who also kindly provided some unpublished LFB data, are acknowledged. Many of the figures were either drafted or resuscitated from earlier versions by Karen Alarcon. TK wishes to thank Richard Arculus for stimulating an interest in the development of the continental crust. This manuscript was prepared while TK held a NERC postdoctoral fellowship, which also supports our ongoing research into the origin of granitic rocks of southeastern Australia. REFERENCES Ablay G. J., Carroll M. R., Palmer M. R., Martı´ J., and Sparks R. S. J. (1998) Basanite-phonolite lineages of the Teide-Pico Viejo volcanic complex, Tenerife, Canary Islands. J. Petrol. 39, 905–936. Albare`de F. (1998) The growth of continental crust. Tectonophysics 296, 1– 14. Allegre C. J., Hart S. R., and Minster J. F. (1983) Chemical structure and evolution of the mantle and continents determined by inversion of Nd and Sr isotopic data: I. theoretical models. Earth Planet Sci. Lett. 66 177-90. Anders E. and Grevesse N. (1989) Abundances of the elements: meteoric and solar. Geochim. Cosmochim. Acta 53, 197 –214. Anderson J.A.C. (1998) Structural, metamorphic, geochemical and isotopic study of metasedimentary enclaves from the southern Lachlan Fold Belt. PhD Thesis, La Trobe University, Melbourne (unpublished). Anderson J. A. C., Williams I. S., Price R. C., and Fleming P. D. (1996) U–Pb zircon ages from the Koetong adamellite: implications for granite genesis and the local basement in NE Victoria. Geol. Soc. Austral. Abstr. 42, 1 –2. Anderson J. A. C., Price R. C., and Fleming P. D. (1998) Structural analysis of metasedimentary enclaves: implications for tectonic evolution and granite petrogenesis in the southern Lachlan Fold Belt, Australia. Geology 26, 119 –122. Anderson J. L. (1983) Proterozoic anorogenic granite plutonism of North America. Geol. Soc. Am. Mem. 161, 133 –154. Anderson J. L. and Morrison J. (1992) The role of anorogenic granites in the Proterozoic crustal development of North America. In Proterozoic Crustal Evolution (ed. K. Condie). Elsevier, Amsterdam, The Netherlands, pp. 263–299. Arculus R. J. (1999) Origins of the continental crust. J. Proc. Roy. Soc. New South Wales 132, 83 –110. Armstrong R. L. (1991) The persistent myth of crustal growth. Austral. J. Earth Sci. 38, 613–630. Atherton M. P. and Petford N. (1993) Generation of sodiumrich magmas from newly underplated basaltic crust. Nature 362, 144–146. Ayres M., Harris N., and Vance D. (1997) Possible constraints on anatectic melt residence times from accessory mineral dissolution rates: an example from the Himalayan leucogranites. Min. Mag. 61, 29– 36. Barbarin B. (1992) Genesis of the two main types of peraluminous granitoids. Geology 24, 294–298. Barbarin B. (1996) Genesis of the two main types of peraluminous granitoids. Geology 24, 295–298. Barbarin B. (1999) A review of the relationships between granitoid types, their origins and their geodynamic environments. Lithos 46, 605 –626.

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Soesoo A. (2000) Fractional crystallization of mantle-derived melts as a mechanism for some I-type granite petrogenesis: an example from the Lachlan Fold Belt, Australia. J. Geol. Soc. London 157, 135– 149. Soesoo A. and Nicholls I. (1999) Mafic rocks spatially associated with Devonian felsic intrusions of the southern Lachlan Fold Belt: a possible mantle contribution to crustal evolution processes. Austral. J. Earth Sci. 46, 725–734. Solar G. S. and Brown M. (2001) Petrogenesis of migmatites in Maine, USA: possible source of peraluminous leucogranite in plutons? J. Petrol. 42, 789–823. Sørensen S. S. (1988) Petrology of amphibolite-facies mafic and ultramafic rocks from the Catalina Schist, southern California: metasomatism and migmatization in a subduction zone metamorphic setting. J. Metamorph. Geol. 6, 405–435. Stein M. and Hofmann A. W. (1994) Mantle plumes and episodic crustal growth. Nature 372, 63–68. Stern C. R. and Kilian R. (1996) Role of the subducted slab, mantle wedge and continental crust in the generation of adakites from the Andean Austral Volcanic Zone. Contrib. Mineral. Petrol. 123, 263–281. Strong B. F. and Hanmer S. K. (1981) The leucogranites of southern Brittany: origin by faulting, frictional heating, fluid flux and fractional melting. Can. Mineral. 19, 163–176. Stump E., White A.J.R., and Borg S.G. (1986) Reconstruction of Australia and Antarctica: evidence from granites and recent mapping. Earth Planet. Sci. Lett. 79, 348–360. Sun S.-S. and McDonough W. F. (1989) Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In Magmatism in the Ocean Basins. Geol. Soc. Spec. Publ. 42 (eds. A. D. Saunders and M. J. Norry) Blackwell Scientific, Oxford, pp. 313–345. Sylvester P. J. (1995) Archean granite plutons. In Archean Crustal Evolution (ed. K. C. Condie). Elsevier, Netherlands, pp. 261–314. Taylor S. R. and McLennan S. M. (1985) The Continental Crust: its Composition and Evolution. Blackwell, Malden, Mass. Taylor S. R. and McLennan S. M. (1995) The geochemical evolution of the continental crust. Rev. Geophys. 33, 241–265. Tepper J. H., Nelson B. K., Bergantz G. W., and Irving A. J. (1993) Petrology of the Chilliwack batholith, North Cascades, Washington: generation of calc-alkaline granitoids by melting of mafic lower crust with variable water fugacity. Contrib. Mineral. Petrol. 113, 333–351. Thompson A. B. (1982) Dehydration melting of pelitic rocks and the generation of H2O-undersaturated granitic liquids. Am. J. Sci. 282, 1567–1595. Thompson A. B. (1999) Some time– space relationships for crustal melting and granitic intrusion at various depths. In Understanding Granites: Integrating New and Classical Techniques. Geol. Soc., Spec. Publ. 168 (eds. A. Castro, C. Fernandez, and J. L. Vigneresse). Geological Society, London, pp. 7–25. Thompson A. B. and Connelly J. A. D. (1995) Melting of the continental crust: some thermal and petrological constraints on anatexis in continental collision zones and other tectonic settings. J. Geophys. Res. 100, 15565–15579. Turner S. P. (1996) Petrogenesis of the late-Delamerian gabbroic complex at Black Hill, South Australia: implications for convective thinning of the lithospheric mantle. Mineral. Petrol. 56, 51–89. Turner S. and Foden J. (1996) Magma mingling in lateDelamerian A-type granites at Mannum, South Australia. Mineral. Petrol. 56, 147–169. Turner S. P., Foden J. D., and Morrison R. S. (1992a) Derivation of some A-type magmas by fractionation of basaltic magma: an example from the Padthaway ridge, South Australia. Lithos 28, 151–179.

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Turner S. P., Sandiford M., and Foden J. (1992b) Some geodynamic and compositional constraints on “postorogenic” magmatism. Geology 20, 931 –934. Turner S. P., Hawkesworth C. J., Rogers N., Bartlett J., Worthington T., and Smith I. (1997) 238U– 230Th disequilibria, magma petrogenesis, and flux rates beneath the depleted Tonga-Kermadec island arc. Geochim. Cosmochim. Acta 61, 4855–4884. Tuttle O. F. and Bowen N. L. (1958) Origin of granite in the light of experimental studies in the system NaAlSi3O8 – KAlSi3O8 –SiO2 –H2O. Geol. Soc. Am. Mem. 74, 153pp. Von Cotta B. (1858) Freiberg, Geologische Fragen, 74p. Vernon R. H. (1983) Restite, xenoliths and microgranitoid enclaves in granites. J. Proc. Roy. Soc. New South Wales 116, 77–103. Vernon R. H. (1984) Microgranitoid enclaves: globules of hybrid magma quenched in a plutonic environment. Nature 304, 438–439. Vernon R. H. (1990) Crystallization and hybridism in microgranitoid enclave magmas: microstructural evidence. J. Geophys. Res. 95, 17849– 17859. Vidal P., Cocherie A., and Le Fort P. (1982) Geochemical investigations of the origin of the Manaslu Leucogranite (Himalaya, Nepal). Geochim. Cosmochim. Acta 46, 1061–1072. Voshage H., Hofmann A. W., Mazzaucchelli M., Rivalenti G., Sinigoi S., Raczek I., and Demarchi G. (1990) Isotopic evidence from the Ivrea Zone for a hybrid lower crust formed by magmatic underplating. Nature 347, 731–736. Waight T. E., Maas R., and Nicholls I. A. (2000) Fingerprinting feldspar phenocrysts using crystal isotopic composition stratigraphy: implications for crystal transfer and magma mingling in S-type granites. Contrib. Mineral. Petrol. 139, 227 –239. Wall V. J., Clemens J. D., and Clarke D. B. (1987) Models for granitoid evolution and source compositions. J. Geol. 95, 731 –749. Wareham C. D., Millar I. L., and Vaughan A. P. M. (1997) The generation of sodic granite magmas, western Palmer Land, Antarctic Peninsula. Contrib. Mineral. Petrol. 128, 81 –96. Watson E. B. and Harrison M. (1983) Zircon saturation revisited: temperature and composition effects in a variety of crustal magma types. Earth Planet. Sci. Lett. 64, 295 –304. Watt G. R. and Harley S. L. (1993) Accessory phase controls on the geochemistry of crustal melts and restites produced during water-undersaturated partial melting. Contrib. Mineral. Petrol. 114, 550–566. Watt G. R., Burns I. M., and Graham G. A. (1996) Chemical characteristics of migmatites: accessory phase distribution and evidence for fast melt segregation rates. Contrib. Mineral. Petrol. 125, 100–111. Wedepohl K. H. (1991) Chemical composition and fractionation of the continental crust. Geol. Rundsch. 80, 207 –223. Whalen J. B. (1985) Geochemistry of an island-arc plutonic suite: the Uasilau-Yau Yau intrusive complex, New Britain PNG. J. Petrol. 26, 603– 632. Whalen J. B. and Currie K. L. (1984) The Topsails Igneous Terrane, Western Newfoundland: evidence for magma mixing. Contrib. Mineral. Petrol. 87, 319 –327. Whalen J. B., Currie K. L., and Chappell B. W. (1987) A-type granites: geochemical characteristics, discrimination and petrogenesis. Contrib. Mineral. Petrol. 95, 407 –419. Whalen J. B., Jenner G. A., Longstaff F. J., Robert F., and Galipey C. (1996) Geochemical and isotopic (O, Nd, Pb, Sr) constraints on A-type granite petrogenesis based on the Topsails igneous suite, Newfoundland Appalachians. J. Petrol. 37, 1463– 1489.

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q 2003, Elsevier Ltd. All rights reserved No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without prior written permission of the Publisher.

Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 349– 410

3.12 Ores in the Earth’s Crust P. A. Candela University of Maryland, College Park, MD, USA 3.12.1 ORE AND CRUSTAL GEOCHEMISTRY 3.12.1.1 Ores, Mineral Deposits, Geochemical Anomalies, and Crustal Composition 3.12.1.2 Physical and Chemical Factors in the Generation of Geochemical Anomalies 3.12.1.3 The Generation of Anomalies 3.12.1.4 The Terrestrial Heat Engine 3.12.1.5 Ore-mineral Solubility 3.12.1.6 Hydrothermal Redistribution of Gold and Antimony: An Example 3.12.1.7 Mineral Deposits 3.12.1.8 Economic Considerations and the Classification of Ore Deposits: The Example of Vanadium–Uranium Deposits 3.12.2 SOME ORE-METAL ANOMALIES FORMED IN THE OCEANIC REALM 3.12.2.1 Seafloor Hydrothermal Sulfide Deposits 3.12.2.2 Sedimentary Exhalative Deposits 3.12.2.3 VMS Deposits 3.12.3 CHROMITE AND RELATED MAGMATIC SEGREGATIONS IN THE OCEANIC AND CONTINENTAL REALM 3.12.4 SOME ORE-METAL ANOMALIES FORMED IN THE CONTINENTAL REALM: THE EXTENDED FAMILY OF (INTERMEDIATE TO FELSIC) MAGMATIC –HYDROTHERMAL SYSTEMS 3.12.4.1 Epithermal Deposits 3.12.5 (GEOCHEMICALLY SIGNIFICANT) PLACERS/RESIDUAL ORES AND CRUSTAL ANOMALIES (ALUMINUM, NICKEL, GOLD) 3.12.6 EPILOGUE REFERENCES

411 411 413 414 414 415 419 421 421 422 422 422 423 425 426 428 429 429 430

mined for profit, they are referred to as ore deposits, and because of the extensive literature developed on the subject, ores will be a central focus of this discussion. Oxygen, silicon, aluminum, iron, calcium, sodium, potassium, and magnesium constitute 99% of the Earth’s crust. If the next four most abundant elements—titanium, hydrogen, manganese, and phosphorus—are included, then nearly 99.9% of the composition of the crust is explained. That is, most of the industrially important elements, aside from iron and aluminum, make up a very small proportion of the Earth’s crust. If we define the scarce elements as those with average abundances less than 1,000 ppm (i.e., elements with abundances less than that of phosphorus), then the 80 or so scarce elements amount to about one-tenth of 1 wt.% (0.1%).

3.12.1 ORE AND CRUSTAL GEOCHEMISTRY 3.12.1.1 Ores, Mineral Deposits, Geochemical Anomalies, and Crustal Composition Variations in the major element composition of the crust are manifested in the proportions of rock-forming minerals. Many of the less common elements can be readily accommodated in these phases or is their commonly associated accessories. In some instances, the less common elements may prove to be anomalous in concentration, and a subset of these anomalies may be of economic interest. Much of the geochemical variation found in the crust conforms to this general motif; in other cases, anomalous concentrations of the less common elements are expressed as the appearance of separate minerals or rock units, defining petrologically distinct geologic entities. When the anomalies can be 411

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The scarce elements form positive anomalies with varying efficiency. To a first order of approximation, the behavior of the scarce elements is largely controlled by the extent to which they differ in their chemical properties from the top ten or so major rock-forming elements, and the extent to which they enter into solid solution in rock-forming minerals (see Chapter 2.09). Mineralogically, the crust (continental plus oceanic) dominantly comprises quartz, feldspar, olivine, pyroxene, mica, amphibole, and clays (here, fine-grained phyllosilicates) with minor oxides, carbonates, sulfides, and other salts. Ronov and Yaroshevsky (1969) performed a norm calculation on the total crust, and suggested that, by weight, the approximate mineral proportions are: feldspar 60 – 65%, quartz 10%, ferromagnesian minerals (olivine, clinopyroxene, orthopyroxene, hornblende, and biotite) 25%, and apatite ,0.5% (a few percent clay and carbonate constituted the balance). These constitute the major rock-forming minerals of the crust of the Earth. However, it is worth noting that the dominant phase in most volcanic rocks is glass. Glass can accommodate a wider range of elements than most crystalline solids, and is also rather reactive in hydrothermal environments. The role of volcanic glass as a source of ore elements has not been extensively explored, but may be an important source of ore metals, as well as chlorine and sulfur. Of the 80 or so scarce elements, many occur in solid solution in major rock-forming minerals (e.g., rubidium, strontium, vanadium, germanium, gallium, scandium, rare earth elements (REEs)), or in solid solution in widely distributed, refractory, or low-solubility accessory phases (e.g., hafnium). Some elements, such as zirconium and phosphorus, form their own minerals. Others, like the base metals (e.g., copper, lead, zinc), semi-metals (arsenic, antimony, bismuth), and

precious metals (gold, silver, platinum group elements (PGEs)) do not possess a high affinity for the minerals that make up the bulk of the crust, and are therefore excluded from all but the most limited solid solution in rock-forming minerals. These elements tend to partition into fluid phases (aqueous and brine phases, or silicate and other naturally occurring melts), and ultimately form their own minerals. In most rocks, these minerals may form small, highly dispersed grains. For example, micron-scale grains of chalcopyrite or cupiferous pyrrhotite may host a high proportion of the copper in a rock, gold may occur as submicron grains of native gold, and gold and arsenic may occur in accessory arsenopyrite (cf. Palenik et al., 2002). If a given volume of the crust is anomalously rich in a given element, but cannot be mined at a profit, we refer to the rock volume as a mineral deposit or a geochemical anomaly. It is tempting to define the term “anomalous” quantitatively as, for example, for a given “number” of standard deviations above the mean. However, most of these definitions are impractical from a geological mining point of view, and a geochemical anomaly is best defined as a significant, positive, geochemical departure from the norm, which may indicate the presence of economic mineralization in a given volume of bedrock. If the elements of interest are concentrated to a sufficient extent to be mined at a profit, then we can call the anomaly an ore. Table 1 lists grades of ores, crustal abundances, Clarke values, and pertinent chemical data for selected ore metals. The Clarke value is the ratio of the concentration of a metal in an ore, relative to its average crustal concentration. These values illustrate the wide range of both crustal concentrations and concentrations in minable ore for these elements. The balance of this paper will examine the nature of the processes that have

Table 1 Data on selected ore metals. Median ore grade for a given element can vary significantly from one deposit type to another. The Clarke value is the ratio of the median grade to the crustal abundance. Element Cu Na Zn As Rb Mo W Pb V Au Ag Ni a

Oxidation state

Deposit type

Gradea (median)

Crustal abundanceb

Clarke value

1 1 2 3 1 4 4 (6) 2 3 0 (1) 0 (1) 2

Porphyry Halite Sedimentary exhalative Sulfide depositsc Lepidolited Climax Skarn Sedimentary exhalative Layered mafic intrusions Veins/Homestake Creed vein Komatiite

0.54% 40% 5.6% ,0.1% Up to 3% 0.19% 0.66% WO3 2.8% ,0.6% ,10 ppm 125 ppm 1.5%

27 ppm 2.3% 72 ppm 2.5 ppm 49 ppm 0.8 ppm 1 ppm 11 ppm 138 ppm 1.3 ppb 56 ppb 59 ppm

200 17 780 ,400 ,610 2,400 6,600 2,500 ,43 ,7,700 2,200 250

All percentages are in weight percent.

b

Continental crust; see Chapter 3.01.

c

Highly variable.

d

In concentrate.

Ore and Crustal Geochemistry produced the heterogeneous crustal distribution of these and other scarce elements. Many of the scarce elements are transported in the crust by complexing with sulfur or chlorine in aqueous solution. Furthermore, many metals are precipitated from aqueous solutions or melts as sulfides, including copper, zinc, lead, molybdenum, rhenium, cobalt, PGE, nickel, cadmium, indium, thallium, antimony, bismuth, arsenic, selenium, tellurium, and mercury. Many of these elements rely on chloride as a complexing agent for aqueous transport. Other elements, such as gold, can be complexed in aqueous solutions by sulfur, but may be precipitated as the native element. Silver, which is not commonly transported as a sulfide complex, may precipitate as the native element or as a sulfide. Sulfur occurs in the continental crust at ,700 ppm (Wedepohl, 1995), and is distributed between reduced and oxidized forms. Hydrothermal solutions commonly contain 0.1 mm to 0.1 m in total reduced sulfur; hence, there is an importance of this element in the transport and deposition of some metals. Chlorine occurs in the crust at concentrations on the order of a few hundred ppm (,300 ppm), and occurs in aqueous solutions over a wide range of concentrations. Chlorine, as well as its heavier congeners, is partitioned into hydrothermal fluids, and ultimately into seawater. Tungsten, uranium, and tin occur in ores almost exclusively as hydrothermally precipitated oxides (Fe, Mn)WO4, UO2, and SnO2, respectively. Most of the ores of these elements are hydrothermal in character. Vanadium, titanium, nickel, and chromium form hydrothermal deposits somewhat less frequently, and are concentrated by crystal settling in mafic intrusions. The noble gases are fugitive, ultimately finding their way into subterranean effluvia or exhalations such as natural gas seeps, volcanic gases, and deep inputs to geothermal systems. Some elements are transitional among groups, usually because of unusual chemical properties. Mercury, for example, concentrates in exhalations, but ultimately can be trapped in organic materials or sulfides as trace cinnabar (HgS). The fluxes of volatile elements in aqueous solutions, as well as the elements transported to sites of deposition by melts, are all affected indirectly by faulting and other elastic manifestations of the upper crust. Therefore, I will devote the next section of this paper to a discussion of fractures in the Earth’s crust, and the engines that drive fluid movement through this porous matrix.

3.12.1.2 Physical and Chemical Factors in the Generation of Geochemical Anomalies Rocks may fail by either fracture or flow. At the low temperatures that obtain at shallow crustal

413

levels, rocks behave elastically and may reach their elastic limit, failing by fracture. With increasing temperature, the resistance of material to flow decreases, and rocks strain by flow before fracturing. With increasing pressure, the elastic strength increases, and so again, rocks will have a tendency to accommodate strain by flow before they reach their elastic limit. For these reasons, rock in the upper-third of the crust behaves differently from deeper rock. In the elastic upper crust, deformation occurs dominantly by frictional sliding on new or pre-existing fractures and by cataclasis, a distributed brittle granulation of rock with progressive decrease in grain size upon deformation. The origin of the stress at a point in the crust can be thought of as a combination of regional (far-field) and localized (near-field) stresses. For example, large-scale regional stresses may be responsible for the shear zones and other fracture networks that allow the upward advection of magmas in active arc zones related to porphyry copper-ore formation. When magma is emplaced at upper levels and continues to evolve physically, it can alter the stress environment through its own thermomechanical evolution (e.g., heating of surrounding rocks, change in volume of the magma due to change in temperature and crystallization, plus volatile exsolution), resulting in a near-field modification of the regional stress state. The diking, fracturing, and concomitant mineralization that forms in and around the crystallizing stock will follow a geometry that is dictated by the time- and space-dependent combination of the near-field and far-field stresses. Generally, these and other upper crustal-fracture networks, comprising new or pre-existing fractures, strongly affect the permeability and porosity of the upper crust, allowing the movement of gases, aqueous solutions, petroleum, and melts. According to Manning and Ingebritsen (2001), geothermal and metamorphic fluid-flux data show that permeability (k, m2) decreases with depth (z, km) in the continental crust according to the equation log k ¼ 214 2 3:2 log z

ð1Þ

above the brittle –ductile transition zone, which crudely corresponds to the upper 10 –15 km of the crust. The brittle – ductile transition zone coincides with log k , 217; which is an effective upper limit on the permeability at which elevated pore-fluid pressures can be sustained. Thus, Manning and Ingebritsen point out that hydrostatic pressure gradients dominate above the brittle – ductile transition zone. When a fracture forms, fluid-filled void held open by tensile forces or by randomly distributed asperities along the walls of the fracture. The fluids may be at below or above the lithostatic pressure. Pressures above lithostatic will almost always cause fracturing: this

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is the case around shallow-cooling magmas, such as those that ultimately yield texturally variable, epizonal granites. This fractured, elastic, permeable rock matrix is the mechanical stage upon which the drama of element redistribution plays out in the upper crust.

3.12.1.3 The Generation of Anomalies The redistribution of chemical elements in the Earth’s crust is effected by a wide range of fluids. The defining characteristic of a fluid, of course, is flow. Flow of a medium, whether a tenuous gas or highly viscous magma, is necessary for the redistribution of elements in the Earth on scales of meters or greater. Vein-type and disseminated hydrothermal ores result from processes including, but not limited to, precipitation from hydrothermal solutions brought about by changes in temperature, pressure, wall – rock reaction, or fluid mixing (that occurs dominantly in a fractured rock matrix). Primarily, magma is transported through the crust by fracture flow. Magma evolution, or differentiation, occurs dominantly by crystal fractionation, and also by a variety of other processes, including assimilation and mingling. One particularly attractive spatial model for differentiation includes the punctuated differentiation model of Marsh (1996) that involves repeated episodes of chambering, fractionation, and removal of liquid (plus some crystals) to higher levels. During these processes, minerals or other phases that are only sparingly soluble in the melt, such as metal oxides, sulfides, PGEs, immiscible oxide and/or sulfide liquids, and saline liquids or vapors may be segregated from the bulk magma by virtue of their significantly different density. Vapor and brine can rise due to their buoyancy. These fluids may be charged with sulfur and chlorine, which can aid in the solubilization of many ore metals, as will be discussed presently. This process appears to be common in high-level intermediate to felsic magmas, and also occurs in mafic systems. In fracture-supported hydrothermal systems, hot fluids rise when surrounded by cooler fluids, or when they are at pressures in excess of the hydrostatic load. Clearly, these and other processes of endogenous element redistribution rely to first order on density differences and upon gradients in the upper crust that are ultimately driven by the irreversible loss of terrestrial heat.

3.12.1.4 The Terrestrial Heat Engine The fundamental engine of the Earth’s endogenous fluid advection system is the irreversible

loss of heat from the interior of the planet to the surface, where it is ultimately radiated into space. The idealized Rayleigh –Bernard convection cell that we can use to model convection of Earth systems can be crudely represented as a Carnot Cycle. At the bottom of the convection cell, matter is heated and consequently expands; the fluid buoyantly rises and cools adiabatically upon expansion, because it has been removed physically from the heat source. In this step, the fluid does “P –V ” (expansion) work, as well as work against gravity. To complete the cycle, the fluid loses heat at a cold (usually upper) boundary, contracts, and consequently sinks. During its downward flow, the fluid is compressed, and also loses gravitational potential energy. The working fluids in “solid – earth” convection systems may be solid mantle, salt (and other sediments), magma, or geothermal fluids; further, upward movement of one fluid (e.g., salt) may be coupled with the downward movement of other “fluids” (e.g., sediment). We can consider the Earth to comprise nested Carnot-convection engines, including the grand convection system of the crust and mantle. The upwelling of magmas and related fluids into the crust can result in the generation of crustally derived magmas. The mantle- and lower crustal-derived magmas are themselves buoyant in the crust, and transport matter to the mid- and upper crust. While in the crust, these magmas can drive the advection and upward transport of magmatic water and locally derived metamorphic water. Further, deep-crustal fluids, as well as more shallowly circulating meteoric fluids are commonly driven to convect, creating large-scale element redistribution systems (see Chapter 3.06 ). Convection may transport crustal fluids and their components reversibly, i.e., with no net transfer of fluid or fluid components from one region to another. Alternatively, net transfer of matter may be effected by changes in phase (e.g., solidification of a magma) or by changes in solubility of some fluid-solute components due to changes in temperature, or pressure, or by chemical reaction along the flow path. For minerals that increase in aqueous solubility with increasing temperature (i.e., those with prograde solubility), there is a net transfer of solute materials from high temperature to low temperature. This is the case for quartz (except for some restricted regions of P – T space) and results in the web of quartz veins that are so common in the upper continental crust. However, some minerals, such as anhydrite, decease in solubility with increasing temperature. This phenomenon, referred to commonly as retrograde solubility, results in the precipitation of minerals in the hotter portions of aqueous convective systems (see also Chapter 3.18). For this reason, oxidized sulfur is precipitated at depth in the oceanic crust.

Ore and Crustal Geochemistry Chemical reaction can fix materials at higher or lower levels. For example, significant concentrations of HCl can occur in magmatic volatiles, and can become reactive upon cooling. Further, magmatic volatiles can also contain SO2, which upon disproportionation yields H2S by the reaction 4H2 O þ 4SO2 ! 3H2 SO4 þ H2 S

ð2Þ

promoting the fixation of metal sulfides in magmatic –hydrothermal systems. The potential for focusing fluid flow around a point heat source such as a magma, with or without significant outward advection of magmatic volatiles, provides for strong lateral (i.e., longitudinal and latitudinal) concentration of ore metals, alteration, and associated chemical and isotopic anomalies. Focusing of geothermal fluids appears to be a critical link in the creation of hydrothermal ores. Unfocused hydrothermal fluid flow produces only geothermal systems, not ore deposits. This focusing is important and is shown by the common creation of large positive geochemical anomalies relative to the rather subdued magnitude of the associated negative anomalies (although in the study of ores as geochemical entities, we must always be cognizant of the inherent economic bias toward the study of not only higher grade and tonnage deposits, but toward positive as opposed to negative anomalies). The presence of cupolas and other apophyses in the roof zones of magma chambers creates thermal point sources that can focus fluid flow. Structures such as releasing bends (Henley and Berger, 2000) act to focus upward surges of magma, forming vertically elongated, high aspect ratio cupolas (Sutherland-Brown, 1976). These conditions obtain most efficiently in local environments of dilation in otherwise compressional regimes. In the New Guinea Fold Belt, at least four major magmatic –hydrothermal Cu – Au (Ag) deposits (Grasberg, Porgera, Ok Tedi, and Frieda River) are located along zones of local dilation. Hill et al. (2002) proposed that, during orogenesis and crustal thickening, strike-slip motion occurred along fracture zones roughly parallel to the convergence vector. Localized zones of dilation opened where these fracture zones intersected major orthogonal faults and other structural discontinuities, facilitating igneous intrusion and subsequent mineralization. The upward advection and subsequent emplacement of magma in a favorable structural environment led to aqueous fluid advection and focusing, and ultimately element redistribution. Finally, most hydrothermal ore formation can be thought of as a type of metamorphism that results in element redistribution. Whereas the hydrothermal alteration commonly associated with hydrothermal anomalies is metasomatic in nature, metamorphism sensu stricto can also be important in

415

ore formation. In the case of greenstone and turbidite hosted mesothermal gold deposits, there is a suggested link between regional metamorphism and the generation of vein-hosted gold mineralization. Metamorphism, together with deformation, can provide a source of fluid, a source of metals, a source of focusing, or all three. The confluence of thermally induced density variations brought about by melting in the lower crust and upper mantle, together with the structural/ tectonic filters operating in the Earth’s brittle upper crust, generally provide the near-point sources of heat and matter required for focusing hydrothermal fluid flow. However, whether fed from underlying magma, metamorphic dewatering, or circulating meteoric water, fluid focusing is not sufficient for the formation of a mineralized system. Hobbs and Ord (1997) point out that fluid focusing is ineffective if the fluid-flow vectors are at low angles relative to isotherms and isobars. Yet, magmas and other fluids may be advected nearly adiabatically into the upper crust. Only when high potential gradients are present, as when fluids are brought into close proximity with the surface, will they flow at high angles to isotherms. Further, magmas at lithostatic pressure may be emplaced into rocks wherein the ambient fluids in fractures are near hydrostatic. Under conditions of strong thermal and mechanical gradients, significant changes in solubility of ore and gangue minerals can occur. Of course, strong gradients in chemical potential can also bring about mineral precipitation. Chemical mixing of different fluids or interaction of fluids with rocks with which they are strongly out of equilibrium result in strong gradients of chemical potential (e.g., the formation of skarns at intrusion/limestone contacts). Generally, the combination of thermally induced density changes in a gravity field, together with strong gradients in thermal, mechanical, and chemical potentials, drive hydrothermal element redistribution in the crust. In the next section of this chapter, I will discuss the general issues related to changes in temperature, pressure, and chemical composition affecting the transport and precipitation of ore metals in hydrothermal systems.

3.12.1.5 Ore-mineral Solubility The overwhelming majority of ore minerals are characterized by very low solubilities in pure water, even at elevated temperatures (Wood and Samson, 1998). Given these low solubilities, it is not clear that pure water could transport sufficient quantities of ore substance to account for the origin of many ores by hydrothermal processes. However, the addition of suitable mineralizers

416

Ores in the Earth’s Crust

(or complexing agents) can raise the concentrations of many ore metals in hydrothermal fluids to levels required for significant transport. Still, some elements are more commonly represented among hydrothermal deposits than others, even after removing the economic bias. It is certainly tempting to assert that the lack of “substantial” solubility is responsible for the dearth of hydrothermal minerals of a given element; however, this assertion is difficult to support upon detailed analysis. In fact, the opposite is more likely the case: the rarity of halite and sylvite as primary minerals in hydrothermal deposits (their abundance as daughter minerals in fluid inclusions notwithstanding) can be attributed to their high solubility, which not only prevents significant deposition, but also leads to their rapid dissolution if deposited. A similar fate may be suffered by other halides. For example, the large cubic chambers in some fluoride-rich pegmatites may represent casts of NaF crystals now dissolved away by subsurface waters (Roedder, 1984). Moreover, gold has a very low aqueous solubility, and is rare, yet hydrothermal deposits abound. The apparent frequency of hydrothermal mineral-deposit anomalies is a combination of competing factors including crustal abundance of the element in question, the magnitude of the changes in solubility over the common ranges of temperature and pressure and chemical potentials of appropriate reactants, the ability of an element to be geochemically camouflaged, as well as the economic bias that may suppress or amplify the apparent abundance of its deposits. The extent to which temperature and pressure affect the solubility of ore minerals is strongly influenced by the chemical composition of the aqueous solutions that may transport the dissolved constituent in question (Wood and Samson, 1998). The set of ligands available for ore-metal complexation are strongly influenced by the general temperature regime. At Earth surface temperatures, biogenic ligands, including carboxylic acids, amino acids, porphyrins, and other polydentate molecules including siderophores and related compounds, as well as ammonia and ammonium, form a wide range of complexes that can enhance equilibrium solubilities, as well as inhibit precipitation kinetically in lowtemperature aqueous environments. More commonly, especially at the higher temperatures and pressures attained in most hydrothermal environments, the dominant ligands are chloride (Cl2), bisulfide (HS2), H2S, hydroxide (OH2), 2 sulfate (SO22 4 ), and bisulfate (HSO4 ). However, oxygen is by far the most common anion in the Earth, and elements such as nitrogen, phosphorus and sulfur, as well as carbon and boron, form common anionic complexes with oxygen. These, in turn, can form ionic compounds in solids, and ions or ion pairs in aqueous solutions.

Although phosphate, sulfate, fluoride, and carbonate are not uncommon anions, their calcium salts (apatite, anhydrite/gypsum, fluorite, and calcite/ aragonite, respectively) are only sparingly soluble, limiting the concentration of these anions in natural waters. Sulfate concentrations in the subsurface are further limited by redox equilibria, fixing some sulfur in sulfide minerals. According to Wood and Samson (1998), hydrothermal solutions range in total reduced sulfur from 0.1 mm to 0.1 m, and sulfate ranges from below detection to a few tenths of a molal. Carbonate is usually in the millimolal to molal range. In contrast, due to the high solubility of chlorides of the major rock-forming elements, naturally occurring aqueous solutions acquire chloride rather readily. Hence, Cl2 is the dominant anion in many natural waters, followed by reduced and oxidized sulfur and (bi)carbonate. Seawater, with a salinity of 3.5 wt.%, has a chloride concentration of 1.9 wt.% (sulfate and carbonate are 0.65 wt.% and 0.14 wt.%, respectively). Seawater is, of course, a major component of ore fluids in the oceanic realm. In contrast, river water is variable in composition, and is, on average, about two and a half orders of magnitude lower in salinity. During base flow, river water is indicative of the composition of shallow groundwater, another important end-member, both thermally and compositionally, for hydrothermal ore fluids. Fluid inclusions in hydrothermal minerals, especially those from granite-related deposits, can range in salinity up to 70 wt.% salt and beyond. Therefore, hydrothermal fluids range widely in salinity, from less than 0.1 wt.% up to hydrous salt melts possessing salinities of many tens of weight percent salt. The salinity of a hydrothermal fluid is a strong determinant of the physical and chemical properties of the fluid. A salt (hydrohalite)-saturated NaCl – H2O solution at 0.1 MPa pressure has a eutectic temperature of 221.1 8C, and a salinity of 23 wt.%; at the same pressure, a saturated solution (,26 wt.%) boils at 108.7 8C. Therefore, even at 0.1 MPa, NaCl can significantly alter the temperature of the phase equilibria of water. Both of these phenomena, freezing point depression and boiling point elevation, are colligative properties, and ultimately are driven by entropic effects that stabilize the mixed aqueous phase relative to the nominally pure solid (ice) and vapor (steam) phase. Two competing effects control the vapor pressure of saturated salt solutions with increasing temperature. Increasing temperature increases the vapor pressure of water; however, a saturated salt solution also increases in salinity with increasing temperature, which tends to decrease the vapor pressure of the solution by dilution of the water component of the liquid. This effect is mostly pronounced for highly soluble salts. In the case

Ore and Crustal Geochemistry

417

the critical point of water, some reversals occur in the solubility of quartz near those temperatures and pressures. Generally, quartz solubility in water varies from a few tenths of a weight percent near granite solidus conditions in the epizone, to a few ppm near Earth-surface conditions. Wood et al. (1987) studied the solubility of the assemblage pyrite, pyrrhotite, magnetite sphalerite, galena, gold, stibnite, bismuthinite, argentite, and molybdenite in aqueous NaCl solutions in the presence of CO2 from 200 8C to 350 8C. They found that relative solubilities followed the sequence Sb . Fe . Zn . Pb . Ag,Mo . Au,Bi in chloride-free solutions, and Fe . Sb . Zn . Pb . Au . Ag,Mo,Bi in chloride-bearing solutions. Further, Wood et al. examined the dependence of ore-metal solubility on the NaCl concentration in the solutions. Their results suggest that iron, gold, and silver are chloride complexed in their chloride-bearing experiments, and that gold and silver are bisulfide complexed in their chloride-free experiments. Few studies show any significant complexation of iron by sulfide complexes, and FeCl2 appears to be the dominant agent of iron transport at elevated temperatures. These results are in general agreement with later studies summarized by Wood and Samson (1998). The same authors, summarizing recent experimental data from a number of sources, state that silver chloride complexes such as AgCl0, 22 AgCl2 2 , and AgCl3 are responsible for aqueous silver transport in a wide variety of hydrothermal environments, although bisulfide complexes may be responsible for silver transport in some

of NaCl, the vapor pressure of salt-saturated solutions is always below the critical pressure at any temperature, allowing for an uninterrupted three-phase equilibrium: halite – vapor – liquid (brine at lower temperatures and hydrous salt melt at higher temperatures, see Figure 1). This three-phase curve is continuous from the four-phase peritectic-invariant point: vapor – hydrohalite – halite – liquid near-room temperature, up to its termination at the pure NaCl vapor –halite– liquid triple point, with a pressure maximum of ,36 MPa near 600 8C. The sizable gap in pressure, ranging over 100 MPa at some temperatures, between the critical pressure and the vapor pressure of the halite-saturated solution at any given temperature (see Figure 1), allows for a significant region of vapor – liquid equilibria over P – T –X space. The applicable temperatures span the full hydrothermal range up to 800 8C, and accessible salinities range from seawater up to pure-salt melts. The applicable subcritical pressures essentially define the Earth’s epizone. The behavior of KCl is very similar to NaCl. The high solubility of both NaCl and KCl in water, together with the relatively high pressures of the critical curves in the binary salt-water systems, allow for extensive aqueous liquid – vapor equilibrium from Earth-surface conditions up to the temperatures and pressures of the granite solidus, and also leads to the importance of chlorides as complexing agents for ore metals in upper crustal fluids. Quartz increases in solubility with increasing temperature and density of the fluid. Because of the variations in density that occur around

1,00 0 ºC

0

Vapor

700

.

Liquid

650

C.C

1.0

0

0 90 0

80

75

85

1,000 80 0

1.5

P (kbar)

600

550 500

0.5 450

H+L+V

400 350

H+V

0 0

10

20

30

40

50

60

70

80

90

100

wt.% NaCl

Figure 1 Isotherms of the NaCl – H2O system illustrating the composition of coexisting vapor (low XNaCl) and brine (high XNaCl) in P – X space. Note that the critical curve (c.c.) is everywhere above the three-phase equilibrium halite – liquid—vapor (source Chou, 1987).

Ores in the Earth’s Crust

cases (Gammons and Williams-Jones, 1997). Further, Wood and Samson (1998) point out that Au(I)-chloride complexes will be important agents of aqueous gold transport at high chloride activity, high temperature, and moderately high oxygen fugacity, with Au(I)-bisulfide complexes occurring over a wide range of conditions. Generally, we might expect that gold-chloride complexes can be important at near and submagmatic conditions, whereas gold-bisulfide complexes will be important at lower temperatures. Wood and Samson point out that the chloride complexes are probably dominant for copper, as suggested by many workers, and the most likely complexes are probably CuCl and CuCl2 2 . Wood and Samson point out that sulfide complexes of copper may play a role, but their stoichiometry and thermodynamics are still a point of contention. They also indicate that tungsten and molybdenum are likely to be transported as tungstate and molybdate species (H2WO4 and H2MoO4, respectively), although systematic data on molybdenite solubility are still lacking. Zinc appears to be transported in hydrothermal environments by a variety of ligands, and Wood and Samson suggest that chloride, bisulfide, hydroxide, carbonate, and bicarbonate complexes all may play roles in hydrothermal zinc transport in the Earth’s crust. Lead appears to be transported as chloride complexes, but carbonate- or sulfidebearing complexes may be important under some conditions. Hemley et al. (1992) studied the solubility of iron, lead, zinc, and copper sulfides in chloride solutions that were rock-buffered in pH and in oxygen and sulfur fugacity, in the range 300 –700 8C; 50 – 200 MPa. Their results show that iron-, copper-, zinc-, and lead-sulfide mineral solubilities decrease with decreasing temperature and decreasing total chloride (see Figure 2), and with increasing pressure. In nature, the HCl concentration (which is the sum of HCl0 and some portion of the ionized Hþ and Cl2) in a hydrothermal solution is controlled by equilibria such as KAl3 Si3 O10 ðOHÞ2ð white micaÞ þ 2KCl þ 6SiO2 ðquartzÞ ! 3KAlSi3 Oð8feldsparÞ þ 2HCl

ð3Þ

for solutions of a given total KCl concentration, pressure, temperature, and salinity (and activity of the specified mineral components). This equilibrium, and related equilibria relating K-feldspar and aluminum silicate minerals such as andalusite, shifts toward higher total HCl with increasing temperature and decreasing pressure. Examination of the reaction CuFeS2 þ 3HCl þ þ 2H2 S

1 2

H2 ! CuCl0 þ FeCl02 ð4Þ

4 Fe

Fe 3 Concentration (log ppm)

418

Zn 2

Cu Pb Pb Zn Cu

1

200

300 400 Temperature (˚C)

500

600

Figure 2 Partly schematic plot of the relationship between solubility of metal sulfides (curves with arrows showing down temperature deposition) and select initial metal concentrations (shown by horizontal lines at temperatures of mineral undersaturation). Drawn for 100 MPa pressure, and for fluids in equilibrium with the quartz –K-feldspar – muscovite buffer with a total chlorinity of 1 m (after Hemley and Hunt (1992)).

shows how ore-mineral solubilities increase as a result of increasing HCl activity. Reactions (3) and (4), taken together, account for the increasing solubilities found for sulfide minerals with decreasing pressure in the experiments of Hemley et al. (1992). In fact, reactions of type (3) generally account for the generation of acidity in otherwise low-acidity hydrothermal fluids as they are heated along prograde paths. Metasedimentary rocks containing these or analogous assemblages can buffer prograding aqueous solutions to higher values of HCl/KCl, HCl/NaCl, (HCl)2/MgCl2 relative to their initial values. Similarly, volatiles exsolved from magmas at shallow levels in the Earth’s crust possess HCl concentrations that ensure that the fluid will possess significant acidity upon cooling, yielding significant metalcarrying capacity and ensuring significant downtemperature hydrothermal alteration upon cooling (Williams et al., 1997; Frank et al., 1999). In a companion paper, Hemley and Hunt (1992) suggest that fluid flow near cooling igneous bodies is nearly adiabatic. Under these conditions, the effects of adiabatic cooling (wherein the work of expansion upon lowering of pressure consumes internal energy, causing temperature to drop) on

Ore and Crustal Geochemistry solubility are counteracted by the positive effect of pressure on ore-mineral solubility, and mineral precipitation does not occur. When the fluids encounter smaller fractures and colder country rock, they will begin to lose heat, and will deviate from the adiabatic path to what Hemley and Hunt call a geothermal path, along which mineral deposition occurs due to cooling, mixing, and wall – rock reaction.

3.12.1.6 Hydrothermal Redistribution of Gold and Antimony: An Example Gold occurs in ores primarily as the native metal. In the crust, where its average concentration is ,3 ppb (Taylor and McLennan, 1995), gold occurs primarily as the native element or in limited solution in other minerals, especially in sulfides and possibly oxides. However, even the gold detected in silicate and oxide minerals may be present as metallic inclusions. A number of workers have suggested that gold can be sequestered in sulfide minerals (Cygan and Candela, 1995; Jugo et al., 1999; Simon et al., 2000). In our laboratory, we have conducted many experiments, wherein gold metal has been equilibrated with silicate melts and aqueous fluids at temperatures ,800 8C and at pressures ,100 MPa (Candela et al., 1996; Frank et al., 2002). We have analyzed coexisting phases by a variety of analytical techniques (Frank et al., 2002; Simon et al., 2002). Our experiments show that the solubility of gold in rhyolitic melts is between 0.3 ppm and 1 ppm at the temperatures, pressures, water, sulfur, and oxygen fugacities that occur in upper-crustal magmatic systems (Frank et al., 2002). Connors et al. (1993) analyzed 129 glassy silicic-volcanic rocks for gold and found that the majority of them (n ¼ 113) contained on the order of 1 ppb. They suggest that granitic systems associated with gold deposits may be elevated in initial gold content, but probably by not more than 4 ppb. Coupling these data with our results on the solubility of gold in a melt at a unit activity of gold, we calculate an activity of gold in those magmas to be on the order of 1023, or ,1,000 times below that required for gold saturation. If this state of gold saturation is typical for crustal magmas, then the source rocks with which crustal melts are last in equilibrium were strongly undersaturated with gold. From that meager evidence, we can suggest that whereas gold is present in ores primarily in native form, in common rocks it is present in a form that significantly reduces its chemical potential, and that form is like solid solution in minerals at the temperatures and pressures of magma generation. If we take crustal melts as a crude indicator of the chemical potential of gold in crustal rocks in general, then

419

from source region to deposition, the ore-forming process must achieve amplifications of gold activity of three orders of magnitude. The state of saturation of crustal fluids with respect to a given ore mineral varies widely from element to element. For example, some high-temperature oil-field brines, such as those associated with gas plays, are near saturation with fluorite, sphalerite, and either metallic lead or galena (Hartog et al., 2001). For gold, the amplification of thermodynamic activity may occur by a variety of processes; however, physical concentration of the fluid as well as efficient precipitation are necessary for the formation of an ore deposit. A number of authors have pointed out the difference in the behavior of gold at low (e.g., , 400 ^ 100 8C), versus high temperatures. In lower-temperature aqueous fluids, gold behaves as a soft (i.e., large and deformable) metal ion, bonding covalently with the soft bisulfide ligand (Wood et al., 1987), whereas at temperatures greater than ,350 – 400 8C, chloride complexes become important. This effect is probably due to the increasingly ionic nature of bonding within aqueous complexes as the mean interatomic distance grows with temperature, in turn leading to the increasing “hardness” of the gold ion and the preference for its interaction with the harder (i.e., smaller and less deformable) chlorine ion. Hence we find that, above ,400 8C, depending upon the pH, total pressure, total Cl2 concentration, and fugacity of H2S (see Gammons and Williams-Jones, 1997), AuCl2 2 is the dominant gold ion (Gammons and Williams-Jones, 1997), with HAuCl02 dominating in very low dielectric fluids (e.g., in near-magmatic vapors) (Frank et al., 2002). At lower temperatures, Au(HS)12 2 dominates (see Table 2). High-temperature goldchloride solutions tend to precipitate gold upon cooling, whereas cooling of lower temperature gold-bisulfide solutions does not lead to gold precipitation. Ilchik and Barton (1997) suggest that precipitation of gold from bisulfide solutions results from ore-fluid dilution or wall – rock alteration. Both dilution and wall –rock reaction can also be important in the deposition of gold at high temperatures. Table 2 Model high-temperature and low-temperature equilibria demonstrating the effects of oxidation and variation of H2S fugacity, as well as the contrasting stoichiometric effects of changing pH, upon gold solubility. 0ð metalÞ þ 2Cl2 High T: AuCl12 2 þ 0:5H2 O ! Au þ þ H þ 0:25O2 þ 0ð metalÞ Low T: AuðHSÞ12 2 þ 0:5H2 O þ H ! Au þ2H2 S þ 0:25O2

420

Ores in the Earth’s Crust

For gold transported at temperatures of less than 400 8C in reduced sulfide-bearing solutions, gold solubility is maximized near the H2S – HS2 –SO22 4 equal predominance point at any given temperature, pressure, and activity of water (see Figure 3). As conditions become more oxidized, the sulfide complexes that stabilize gold in aqueous solution break down, and gold metal precipitates. As pH increases (in the HS2 predominance field), the redox reaction 0:5H2 þ Auþ ! Au0ð metalÞ þ Hþ

ð5Þ

proceeds to the right, causing gold precipitation. The net equilibrium at gold saturation under these conditions is 0ð metalÞ AuðHSÞ2 þ 2HS2 2 þ 0:5H2 O ! Au

þ Hþ þ 0:25O2

ð6Þ

(where the reaction is written involving oxygen and water rather than hydrogen and water so as to show the effect of oxygen fugacity). Note that for a given water activity or fugacity at a given temperature and pressure, the oxygen and hydrogen fugacities vary inversely. If pH decreases into the H2S-dominant field, the hydrogen –gold redox reaction is coupled with the conversion of HS2 to H2S, yielding gold precipitation with decreasing pH as illustrated by the low-temperature mass action expression shown in Figure 2. Hence, depending upon whether gold is being transported in the field of H2S or HS2 predominance under –30

–8 HSO–4 SO2– 4

–36

–6

Hem Au(HS)–2

2

–2

0

Py –48

Log ƒH

2

Log ƒO

pm 10 p

ppb 100

b 1 pp

–42

–4

Mag

Po

2

H2S0

HS–

–54 0

2

4

6

8

10

12

pH

Figure 3 log fO2 – pH diagram for the solubility of gold as Au(HS)2 2 , at 200 8C and saturated water vapor pressure. Solid lines delineate mineral-stability fields; dotted lines delineate the fields of dominance for dissolved sulfur species (total dissolved sulfur ¼ 0.01 m), and dashed lines show gold solubility contours. Drawn for an activity of water equal to unity (after Wood and Samson, 1998).

reduced conditions, gold may be precipitated by either pH decrease or pH increase, respectively, as long as the oxidation state of the solution remains in the reduced-sulfur field of predominance. Reduction can also precipitate gold, as is shown by both the gold-bisulfide-bearing equilibria, as long as the oxidation state of the solution remains in the reduced sulfur field of predominance. Only when a sulfide solution is oxidized to sulfate can gold precipitate from an aqueous solution by oxidation. This reaction can be written as 15 AuðHSÞ2 2 þ 0:5H2 O þ 4 O2 þ ! Au0ð metalÞ þ 2SO22 4 þ 3H

ð7Þ

Williams-Jones and Normand (1997) suggest that the conditions of maximum gold concentration in hydrothermal fluids correspond closely to the conditions of maximum antimony concentration, where the dominant aqueous antimony species is HSb2S2 4 . Because stibnite (Sb2S3) solubility increases rapidly with temperature, and because antimony is a rare element, the ratio of the concentration of antimony in hydrothermal solutions to its solubility limit is usually rather low, until temperature drops to the general range of 150– 300 8C. Under the restricted conditions of maximum gold solubility, stibnite precipitation is considered by Williams-Jones and Normand to be controlled by the mass action expression þ HSb2 S2 4 þ H ! Sb2 S3 þ H2 S

ð8Þ

suggesting that acidification can cause stibnite precipitation. Williams-Jones and Normand point out that many of the deposits exhibiting a strong antimony – gold association occur in settings which provide for an alkaline environment of transport (talc-carbonate or limestone host rocks), followed by precipitation in an environment capable of aqueous acidification, such as in fractures hosted by phyllosilicate-rich rocks, wherein reactions such as Equation (3) can produce hydrogen ions, promoting both stibnite and gold deposition. Fluids that are more acidic and oxidized can carry antimony as Sb(OH)03, precipitating stibnite upon reduction if sufficient sulfur is available, hence the occurrence of many black shale-hosted stibnite deposits. These ore fluids, however, could not transport much gold because of their oxidized nature, unless the transport temperatures were high enough to promote significant chloride complexing of gold. Antimony is characteristic of a class of elements that possess a combination of high solubility at high temperatures, strongly decreasing solubility with decreasing temperature and/or pressure, and low abundance. These characteristics tend to restrict this class of elements, which includes selenium, tellurium, mercury, thallium as well as antimony, and sometimes arsenic and bismuth,

Ore and Crustal Geochemistry to low-temperature deposits and in many cases to igneous-related deposits.

3.12.1.7 Mineral Deposits The nonlinear interactions in the crust among thermal gradients, mechanical properties of rocks in shallow crustal environments, and irreversible chemical reactions, especially those operating in strong gradients with significant entropy production (dissipative environments), lead to the uneven, power-law, or lognormal distribution of element concentrations in the crust, geochemical anomalies, mineral deposits, and ores (Shen and Zhao, 2001). With regard to ore systems, we refer to the average concentration (by weight) of the metal in an arbitrarily chosen mass of rock as the grade, usually expressed in percent, or ppm (g t21). The estimated mass of rock that is characterized by that average grade is referred to as the tonnage. Grade and tonnage data were originally used by Lasky (1950) to show the inverse relationship between average grade and the logarithm of the cumulative tonnage of porphyry copper deposits. For comparison, the Earth’s continental crust has a volume ,1010 km3, with the volume of the upperone kilometer occupying a little over 108 km3. Using the upper-one kilometer as a standard of comparison (because mining is limited to the upper few kilometers of the Earth’s surface), a very large ore deposit with a “tonnage” or mass on the order of 3,000 Mt (e.g., a world class porphyry copper in the top 5%) and with a volume on the order of 1 km3 would occupy one part in 108 of the upper onekilometer-thick shell of the Earth’s land surface (a millionth of a percent). Of course, most ore or mineral deposits are much smaller. If we consider the upper ten kilometers of the Earth’s crust rather than just the upper one kilometer, the proportion would drop by an order of magnitude. In terms of the mass proportion of the continental crust, the same giant ore deposit would represent ,0.1 ppb by mass. Clearly, just as mining occupies a vanishingly small proportion of the area of the Earth’s surface, ore deposits do not contribute significantly to the bulk inventories of any element. However, whereas ores and related anomalies may not be quantitatively important, they are important as examples of the limits of enrichment that are possible for any given element.

3.12.1.8 Economic Considerations and the Classification of Ore Deposits: The Example of Vanadium–Uranium Deposits Under the oxidation state conditions that prevail in the crust (a range of a few log units straddling the QFM-NNO region of log fO2 – P – T space),

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vanadium occurs dominantly in the trivalent state. The ionic radius of V3þ is virtually identical to that of ferric iron, so it is easily accommodated by ferric iron-bearing minerals. There is roughly 400 –500 times more iron than vanadium in the Earth’s crust. That is, iron-bearing minerals act as “solvents” for trivalent vanadium, and the bulk of the crust’s inventory of vanadium resides in ferromagnesian rock-forming minerals. Economic sources of vanadium are commonly considered to be comparatively rare, and include vanadiferous magnetite, by-product vanadium from uranium mining, and recovery from petroleum residues; other sources include vanadium-bearing iron slag, and flyash. The chemical weathering of vanadium-bearing minerals, the release of trivalent vanadium and its subsequent oxidation, lead to the transport and deposition of vanadates in near-surface environments, where pentavalent vanadium minerals precipitate. Much is made in the economic geology literature of the fact that vanadium is a by-product of uranium mining; however, this result is not so much geochemical, as economic. A similar relationship exists between silver and gold in many deposits. As this is a Treatise on Geochemistry, and the subject is ores, it will be instructive to examine how the economic bias in ore deposit information must be considered in a geochemical context. To discuss this issue further, some background in uranium – vanadium deposits is in order. Sandstone hosted uranium –vanadium deposits, which account for ,20% of the world’s uranium, are found in medium to coarse-grained sandstones deposited in continental-fluvial or marginalmarine sedimentary environments. These units acted as aquifers for the mineralizing oxidized aqueous solutions, and are frequently bounded above and below by lower permeability horizons that act as aquacludes. Uranium, vanadium, and other redox sensitive elements precipitate upon encountering reducing conditions that are produced within the sandstone by carbonaceous biogenic material, aqueous or solid sulfide, or other reduced substances. Three main types of sandstone deposits have been recognized: rollfront deposits, which are characterized by arcuate discordant uranium – vanadium mineralized zones; tabular deposits, with irregular, elongate, tabular mineralized zones concordant to the bedding; and structurally controlled deposits that occur in sandstones adjacent to permeable fault zones. These ore bodies are commonly low to medium grade (,0.1% U3O8) and their size ranges up to 50 kt of U3O8. According to Fitch (1997), deposits of carnotite on the Colorado Plateau were the first major sources of radium from 1912 to 1922 and of vanadium from 1924 to 1945. Uranium was discarded as part of the vanadium mine tailings, and ultimately, these tailings became a source for

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Ores in the Earth’s Crust

much of the uranium used for atomic weapons. Mining of uranium ores for military use began in 1947 and continued until 1970. The use of uranium for nuclear power plants began in the early 1950s (Fitch, 1997). As a result of the importance of uranium since 1945, what were once referred to as vanadium deposits became “uranium,” or “uranium – vanadium” deposits. In fact, most tabular sandstone uranium deposits in the USA contain more vanadium than uranium, commonly 7– 12 times as much; i.e., geochemically, uranium was a by-product of vanadium mining, even though from 1945 up to the time of the Three Mile Island/Chernobyl incidents, these deposits, economically, were uranium mines. The greatly increased demand for vanadium in the late 1990s (spurred in part by the increased demand for steel, worldwide) was demonstrated by the sharp rise in the price of vanadium from $3.30 per pound of V2O5 in January 1997 to $6.70 in January 1998; however, the price of U3O8 was still $12 per pound at that time; at that point, many tabular sandstone uranium deposits became uranium – vanadium deposits. Therefore, we must be careful how we interpret ore-related information (e.g., vanadium is “only” a by-product of uranium mining) from a geochemical framework. The next section discusses select deposit types from a variety of crustal settings and shows how density, temperature, pressure, and chemical potential gradients play a role in the genesis of ores by hydrothermal, igneous, or surficial processes.

3.12.2 SOME ORE-METAL ANOMALIES FORMED IN THE OCEANIC REALM 3.12.2.1 Seafloor Hydrothermal Sulfide Deposits Seafloor hydrothermal deposits include sedimentary exhalative deposits (Sedex), volcanogenic massive sulfide (VMS), and Besshi massive sulfide deposits. Compared to VMS deposits, Sedex deposits are fewer in number but generally an order of magnitude larger (Goodfellow, 2000). Sedex deposits are major sources of zinc and lead, and can be considered as a subgroup of massive sulfide deposits associated with sedimentary rocks. Both the sediment and volcanogenically hosted massive sulfide deposits are characterized by lenses or sheets of ore that comprise at least 60% sulfide, hence the origin of the name. As such, these deposits represent major crustal anomalies of sulfur. 3.12.2.2 Sedimentary Exhalative Deposits Sedex deposits constitute many of the world’s great ore deposits, including the Sullivan mine in Canada, Mount Isa, the Broken Hill protolith, and

HYC in Australia, and Rammelsberg in Germany. The ores are characterized by beds and laminations of sulfides that commonly comprise sphalerite, galena, pyrite and pyrrhotite possibly with chalcopyrite, and barite. The lenses or layers of ore range from centimeters to tens of meters in thickness. Multiple layers may occur over a vertical interval that can be on the order of a kilometer, and can extend up to several kilometers along strike. The term Sedex, or “sedimentary exhalative,” is a generic name that reflects the current understanding of the genesis of these deposits as having formed by the precipitation of sulfides from hydrothermal fluids vented or “exhaled” on to the seafloor. According to MacIntyre (1995), the major Sedex metallogenic events are Middle Proterozoic, Early Cambrian, Early Silurian, and Middle to Late Devonian to Mississippian in age. The MiddleProterozoic and Devonian– Mississippian events are recognized worldwide. Briskey (1986) gives the median tonnage for zinc –lead– gold deposits as 15 Mt, with 10% of deposits in excess of 130 Mt. According to MacIntyre (1995) the median grades worldwide are zinc 5.6%, lead 2.8%, and silver 30 g t21. The Sullivan deposit, one of the supergiant Sedex deposits, has a total tonnage of 155 Mt with 5.7% zinc, 6.6% lead, and 7 ppm silver. The lead concentration commonly correlates with the silver concentration (Menzie and Mosier, 1986). Lateral Cu ! Pb ! Zn ! Ba zoning sequences extend outward from vent or feeder zones; ammonium anomalies may also be present. Sedex deposits generally form in restricted sedimentary basins or half grabens bounded by syn-sedimentary growth faults. The exhalative vents generally occur along the faults and the exhaled brine accumulates in adjacent seafloor depressions. As Sangster (2002) points out, most Sedex deposits are vent-distal. He proposes that saline fluids discharged onto the seafloor and flowed away from the vent as “bottom-hugging” fluids. He suggests that their high densities and velocities prevented them from mixing with overlying seawater, and precluded significant cooling and dilution of the ore fluid until they came to rest in a seafloor topographic low. There is a close spatial relationship between Sedex deposits and mafic volcanic rocks. Compared to VMS deposits, however, Sedex deposits are typically lead – zinc rich and copper poor, and formed under lower temperature conditions (i.e., less than 260 8C). The redox state of the mineralized brines, which in turn is reflected in whether the ore fluid is dominated by sulfate (oxidized) or sulfide (reduced), is important for controlling minor element associations in Sedex deposits. Cooke et al. (2000) point out that weakly acidic to weakly alkaline oxidized brines can precipitate siderite

Some Ore-metal Anomalies Formed in the Oceanic Realm but are incapable of carrying significant gold, tin or barium in solution, and that therefore McArthur River-type deposits do not contain anomalous concentrations of these elements. Reduced, acid brines can carry high concentrations of barium or tin; if these reduced brines contained sufficient sulfide, the deposits may contain anomalous or ore-grade gold (Cooke et al., 2000). Across the SO22 4 /H2S boundary, the solubility of lead and zinc increases rapidly with increasing oxygen fugacity, due to the decreasing activity of reduced sulfur (Large et al., 1998). Temperature decrease and fluid mixing, addition of H2S, and pH increase can all be effective processes for the deposition of zinc and lead sulfides from reduced brines. In contrast, sulfate reduction and/or addition of H2S (by way of fluid mixing or interaction with earlier formed sulfides) may be an important process for the deposition of galena and sphalerite from oxidized brines (Cooke et al., 2000). In some cases, metalliferous hydrothermal fluids may vent into anoxic basins, precipitating a different suite of ore metals, forming, for example, deposits such as the nickel –molybdenum –zinc – PGE-bearing black shales of China (Murowchick et al., 1994). These deposits comprise thin layers or other segregations of sulfide minerals in blackshale sequences. The geochemical anomalies are marked by above-average concentrations of nickel, molybdenum, rhenium, zinc, gold, PGE, organic matter, phosphates (usually as discrete beds), barium, selenium, arsenic, vanadium, uranium, and sulfur (see Table 3).

3.12.2.3 VMS Deposits VMS deposits are lens-like to sheet-like volumes of sulfide mineral-rich rock bodies

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associated with volcanic rocks. Generally, the deposits form in deep-sea environments and belong to the larger class of concordant massive sulfide deposits, which includes all massive or semi-massive sulfide deposits formed by the discharge of hydrothermal solutions in nearseafloor environments. They range in age from ,3,500 Ma for deposits of the Pilbara Craton to the modern black smoker deposits of the midocean ridges. The deposits occur at mid-ocean ridge spreading centers, in the fore-arc areas of convergent zones, back-arc basins, and in subaqueous continental rift zones. According to Franklin et al. (1981), ,80% of the Earth’s VMS deposits are in arc-related successions. The major metals in these anomalies include copper, zinc, and lead, as well as gold, silver, and tin. Generally, the Zn : Cu ratio increases upward and outward from the core of the massive sulfide lens. In this chapter, VMS deposits are separated into three classes: the Cyprus-type, Noranda/ Kuroko-type, and Besshi-type deposits. Cyprus-type VMS deposits contain small, stratiform, basalt-hosted, medium-grade ores dominated by copper, and zinc as the prime commodities. The Upper Cretaceous Troodos ophiolite complex in Cyprus is the type locale for this class of VMS deposit. Generally, these deposits take the form of lens-like masses of pyrite, characteristically associated with basaltic pillow lavas and cherts. The related sequence of mafic and ultramafic rocks associated with Cyprus-type VMS deposits are generated at mid-ocean ridge (MOR) or back-arc spreading centers. Cathles (1981) summarizes a number of studies that suggest that these local redistributions of copper, zinc, iron, sulfur, and other metals occur on oceanic crust that is significantly less than 1 Myr old. The geochemical anomalies

Table 3 Summary of tectonic environments and ore components in selected deposit types. Deposit type Sedex

Tectonic environment

Primary ore metals

Other metals present

Fault-controlled basins in intracratonic or continental margin settings Ophiolite complexes

Zn, Pb, Ag

Cu, Ba, Sn, Au

Cu Cu, Pb, Zn, Ag, Au

Porphyry deposits

Local extensional areas of arc settings Extensional environments in the oceanic realm Arcs

Magmatic Ni – S Epithermal

Mid-continent rifts Arcs

Cu, Ni, PGE, Co Cu, Au, Ag

Ni – Mo – Zn – PGE-bearing black shale deposits

Continental platforms

Ni, Mo

S, Fe, Cu, Au, Ag, Zn, Co, Cd, Pb S, Fe, Ba, As, Se, Cd, Sn, Bi, B S, Fe, Pb, Ni, Mn, As, Sb, Sn, Mo, B Re, S, Fe, As, Se, Bi, W, B, Sr, Zn, Pb, Co, V, PGE, Sn S, Fe As, Sb, Hg, Pb, Zn, Cu, Ba, F, Mn, Mo, Se Zn, Pt, Pd, Au, P, Ba, Zn, Se, As, S, U, V

VMS-Cyprus VMS-Noranda/Kuroko Besshi

Cu, Zn, Ag, Co Cu, Mo, Au

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Ores in the Earth’s Crust

comprise sulfur, iron, copper, gold, silver, zinc, cobalt, cadmium, and lead. The main ore zones may be underlain by a pipe-shaped stockwork of copper-rich “stringer-zones” consisting dominantly of quartz-sulfide (and in some cases paragonite-bearing) veins in heavily chloritized basalt. Ore mineralogy includes pyrite þ chalcopyrite þ magnetite þ sphalerite, with minor galena and pyrrhotite. The main ore zones are structurally controlled by normal faults that served to focus fluid discharge onto the seafloor. Contemporary black smoker deposits found at spreading ridges appear to be modern-day analogues of Cyprus-type VMS deposits. The Noranda and Kuroko types of VMS deposit represent seafloor hydrothermal anomalies of sulfur, iron, copper, zinc, lead (referred to sometimes as copper – zinc ^ lead massive sulfide deposits, because these are dominant economic commodities) which form in what are essentially intermediate to felsicpyroclastic domes in submarine calderas. The Noranda (Kidd Creek) type deposits occur in Archean– Proterozoic greenstone belts and are lead poor. These are to be distinguished from the Cyprus-type VMS deposits that are also lead poor, but which form at MOR spreading centers (copper – zinc massive sulfide deposits), and are associated with mafic rocks. About half of all VMS deposits are associated with intermediate to felsic-volcanic rocks; the balance are associated with mafic rocks, or mixed-volcanic/sedimentary successions. The Kuroko-type deposits, which are associated with intermediate to felsicvolcanic rocks, are stratabound, and are usually stratiform. The host rocks include rhyolite, dacite, and minor basalt. Fine-grained clastic sediments or sedimentary rocks may also be present. Ore mineralogy generally consists of an upper zone of “black ore” which is dominantly pyrite þ sphalerite þ chalcopyrite ^ pyrrhotite ^ galena ^ barite, and a lower zone of “yellow ore” which is dominantly pyrite þ chalcopyrite ^ sphalerite ^ pyrrhotite ^ magnetite. Barite and chert are common gangue minerals. The conical stringer vein or stockwork feeder zone generally contains pyrite þ chalcopyrite. Gold and silver, as well as zinc, lead, and barium tend to increase toward the more distal portions of the anomaly. Black chlorite and talc are common alteration minerals. The anomalies usually comprise sulfur, iron, copper, zinc, lead, barium, arsenic, silver, gold, selenium, cadmium, tin, bismuth, and boron. Besshi deposits comprise thin sheets of massive, well-layered pyrrhotite þ chalcopyrite þ sphalerite þ pyrite with minor galena and cobalt minerals interlayered with terrigenous clastic sedimentary rocks and calc-alkaline

basaltic to andesitic volcanic rocks. Ultramafic material can be present. Their grade and tonnage are generally lower than those of other massive sulfides, but cobalt may be present as an additional commodity. Sulfur, iron, copper, zinc, lead, cobalt, nickel and manganese, arsenic, antimony, tin, molybdenum, and boron may be present in the anomaly. As in many other ore deposits, Besshi anomalies can have high-aspect ratio layers that are just meters thick, yet extend laterally on the kilometer scale. The rather eclectic mixture of metals probably reflects the mixed sources that are present near Besshitype hydrothermal systems. Copper and zinc are typical products of the alteration of mafic volcanic rocks, whereas significant cobalt and nickel may represent an ultramafic component (Candela et al., 1989). However, tin, lead, molybdenum, and boron are more characteristic of continental crust-related sources. In the formation of a massive sulfide deposit, cold seawater descends along fractures in the far recharge regions of the seafloor hydrothermal system, is heated, and leaches metals and sulfur from the volcanic succession. The fluid can be drawn into the central regions of the convecting system that is focused on the magmatic heat source. The seawater convection cell is driven by the irreversible loss of heat from cooling submarine igneous bodies. Boiling is suppressed during the formation of VMS deposits, allowing the ore fluid to exit onto the seafloor; for this reason, they must form at or above a minimum pressure for any given temperature of the hydrothermal fluid. The minimum pressure to prevent boiling increases as the fluid temperature increases. For the most efficient ore formation in either the Kuroko – Noranda or Cyprus-type VMS deposits, discharge flow is best controlled by a few large fractures, which focus flow through a narrow lateral region. However, recharge favors ore formation if it is diffuse, with long transit times. Here, large through-going fractures would short-circuit the leaching path, reducing the contact between the heating fluid and the rock with which it is strongly out of chemical equilibrium (Cathles, 1981). The fluid that exits onto the seafloor at a submarine hydrothermal vent is saline, reduced, acidic, metalliferous, and sulfidic. During prograde hydrothermal alteration, calcium, sodium, and some iron are leached from the basalt, along with ore metals and sulfur, whereas seawater magnesium is fixed in chlorite and related minerals. Seawater sulfate is reduced by ferrous iron in glass and silicates (basalts are commonly half glass), producing ferric iron and H2S, with the oxidized iron fixed in magnetite, hematite, or silicates. HCl is also produced by magnesium fixation, as illustrated by the

Chromite and Related Magmatic Segregations schematic reaction 4H2 O þ 2Mg2þð swÞ þ CaAl2 Si2 Oð8glass;fsÞ chlÞ ! Mg3 Si2 O5 ðOHÞð4chlÞ þ Al2 Si21 Mgð21

þ SiOð2aqÞ þ Ca2þ þ 4Hþ

ð9Þ

and by other reactions. This makes the fluids more aggressive with respect to leaching of metals (manganese, iron, copper, barium, zinc, etc.) at higher temperatures. The mixing of these waters with bottom ocean water near 2 8C causes mineral precipitation, dominantly by cooling upon dilution. Cathles (1981) points out that the salinity of the discharged fluid also determines its fate. Brines are ponded in depressions in the seafloor, as has occurred in the past in the Red Sea. The hydrothermal fluids may interact with the oceanic crust at any temperature between the temperature of ambient seawater and temperatures on the order of the solidus of basaltic magma (,1,000 8C). However, the exit temperatures on the seafloor in modern hydrothermal vent areas can be on the order of 350 –400 8C. Although significant permeability probably does not develop until the rock cools and cracks (by differential thermal contraction) at or near 800 8C (Lister, 1974), other more subtle thermodynamic factors may be involved in controlling the temperature seafloor hydrothermal fluids. Jupp (2000) suggests that this discrepancy between rock and fluid temperatures is due to the variations in the thermodynamic and transport properties of water near the critical point which impose an upper limit of ,400 8C on the temperature of black smokers. Jupp (2000) predicts that temperatures greater than ,400 8C are restricted to a thin boundary layer near the base of the region that supports the aqueous convection cell and which may correspond to the reaction zone observed in ophiolite complexes. He refers to fluids near the critical points as having maximum “flexibility.” Thus, the temperature of venting of the fluids (350 – 400 8C) appears to be related to the temperatures at which seafloor vent fluids experience maxima in those physical properties that promote flow of the fluid in a convective regime.

3.12.3 CHROMITE AND RELATED MAGMATIC SEGREGATIONS IN THE OCEANIC AND CONTINENTAL REALM Most of the world’s chromium reserves occur in large, layered mafic/ultramafic intrusions such as the 2.06 Gyr old Bushveld Complex in the Republic of South Africa, where nearly monomineralic layers of chromite appear to result from the mixing of primitive and evolved liquids that

425

lie on or near the chromite/olivine cotectic. However, significant commercial production of chromium also results from ophiolite-hosted podiform chromite or chrome-spinel deposits which occur in alpine-type ultramafic bodies that have been incorporated into the continental crust by obduction. In these deposits, the chromite occurs as podiform bodies found within harzburgite– dunite lithologies. The dunites and harzburgite represent cumulates from the crystallization of MOR basalt, rather than depleted mantle material (otherwise, the banded chromitites are hard to explain). Examples include the deposits of Cyprus, Oman, Newfoundland, and minor showings in many ancient suture zones such as the Appalachians of eastern North America (cf. Wylie and Candela, 1999). The large layered mafic intrusions are gabbroic, and were emplaced into stable cratons, with the economically important bodies all Proterozoic in age. The intrusions are generally in the form of large upward flaring funnel-shaped bodies, thousands of meters thick, and cover thousands of square kilometers. The uniting factor that may link the chromite deposits of the layered mafic intrusions with the podiform-type deposits is their association with rifting, the former continental (and failed), and the latter oceanic. Ultimately, chromium is concentrated due to the highly refractory (insoluble) nature of chromite in silicate melt systems, which causes chromite to precipitate early from mafic magmas. Irvine (1977) suggested that nearly monomineralic layers of chromite could form by contamination of the magma by felsic material or by mixing of more primitive magma with more evolved residua. Mixing can induce precipitation of chromite only because of the curvature of the chromite/olivine cotectic. Curvature is common for cotectics of sparingly soluble minerals in melts (e.g., chromite, sulfides, and apatite) especially when those minerals differ significantly in composition from the bulk melt composition; curvature of the cotectic crudely reflects changes in solubility of the mineral in the silicate melt with temperature. In general, contamination of less-evolved melt with more evolved material may cause saturation of the melt with respect to chromite (or sulfide, apatite, or other minerals that behave similarly). PGE sulfides may also crystallize as suggested by Kruger et al. (2002) for the Bushveld. The PGEs are found associated with minor sulfide-rich bands comprising dominantly pyrrhotite, pyrite, and pentlandite, along with minor but important platinum sulfides. The sulfides are interstitial to the chromite and other minerals of the chromitites. Significant vanadium also occurs in separate magnetite layers within the Bushveld system, representing yet another resource derived from

426

Ores in the Earth’s Crust

gravitationally segregated precipitates from mafic magmas in rift environments. Primary nickel sulfide deposits (with associated copper, cobalt, and PGE mineralization) are typically associated with plutonic and volcanic rocks of mafic or ultramafic composition. For example, the Voisey’s Bay nickel –copper –cobalt sulfide deposit of Labrador is hosted by the troctolitic Voisey’s Bay intrusion, a member of the Proterozoic Nain Plutonic Suite. As of 1999, the total resource was estimated to be 137 Mt of ore containing 1.6% nickel, 0.85% copper, and 0.09% cobalt. Contamination of the mafic magma by a sulfide-bearing paragneiss promoted the development of sulfide liquids, which interacted with successive magma batches resulting in ore-metal enrichment in the sulfide phase (Kerr, 2000). The saturation of a silicate melt with an immiscible sulfide liquid is the fundamental process that forms magmatic sulfide deposits. Droplets of sulfide liquid settle through less dense silicate magma; cobalt, copper, nickel, and PGE partition strongly into the sulfide liquid phase at levels 10– 100,000 times those in silicate liquids (Foose et al., 1995). The mineralized zones at Voisey’s Bay are associated with fragment-bearing troctolites and olivine gabbros related to the conduit dikes, rather than as basal accumulations at the floor of the chambers (Evans-Lamswood et al., 2002). In general then, magmatic segregations of oxides, phosphates (nelsonites?), or sulfides can be expected during contamination of the magma, or by mixing of early and late crystallization products. They generally result from the upward advection of magma, contamination of the magma, precipitation of minerals due to a strong decrease in solubility upon compositional evolution of the melt (e.g., chromite and phosphate), concentration of elements by partitioning from large masses of magma into small masses of “getter” phases (e.g., sulfides in the nickel-sulfide deposits), and ultimately, to gravitational density-driven segregation.

3.12.4 SOME ORE-METAL ANOMALIES FORMED IN THE CONTINENTAL REALM: THE EXTENDED FAMILY OF (INTERMEDIATE TO FELSIC) MAGMATIC–HYDROTHERMAL SYSTEMS Over 65% of all mined copper comes from deposits associated with igneous rocks, i.e., the porphyry, skarn, epithermal, replacement, and massive sulfide deposits, with just over half coming from porphyry type ores (Singer, 1995). The next most important category for contained copper, according to Singer, is sediment-hosted deposits (23%). In contrast, the proportion of zinc

from igneous-related systems is half that of copper. Singer calculates that only 19% of mined gold comes from igneous-related systems, and about one-third of that is in epithermal deposits. Sixty percent of all discovered gold is estimated by Singer to come from modern placers and residual deposits, and ancient placers such as the Witwatersrand deposits in South Africa. Molybdenum is even more strongly associated with igneous rocks than is copper, with well over 99% of the world’s molybdenum coming from porphyry type deposits (Kirkham and Sinclair, 1996). Porphyry deposits are the archetype of the magmatic –hydrothermal class of deposits (also known more descriptively as granite-related deposits). They exhibit a strong spatial and temporal relationship between epizonal, porphyritic/ variably-textured intermediate to felsic igneous rocks and usually steep-walled, crudely cylindrical sulfide-rich ore bodies. The high-level igneous stocks associated with porphyry and related deposits typically form, as mentioned earlier, the roof zones of shallow plutons where vertically elongated, high-aspect ratio cupolas (Sutherland-Brown, 1976) act to focus upward surges of magma and associated magmatic volatiles. Porphyry copper ore occurs as disseminations or stockwork veins yielding grades of copper on the order of 0.4– 1%, with subordinate molybdenum and gold. Porphyry molybdenum deposits have grades of a few tenths of a percent MoS2, and porphyry gold deposits contain on the order of 1 ppm gold. Characteristic alteration styles are present which represent the combined effects of fluid cooling and the mixing of magmatic and meteoric waters. The formation of granite-related ore deposits is a by-product of the irreversible transfer of magmatic heat and mass from the Earth’s interior to its surface. Magma emplacement at epizonal levels leads to magmatic volatile phase (MVP) saturation if the system is initially volatile-phase undersaturated, or to further exsolution of volatiles if the magma is already volatile saturated (Candela, 1989). The MVP is thought to be a critical agent in ore formation because of the affinity displayed by many ore metals for a chloride-bearing volatile phase relative to rockforming minerals (Candela and Piccoli, 1995), and because of the high fluidity and the buoyancy of the MVP in the magma. These physical attributes may allow the volatile phase to accumulate near and above suitable apical magmatic structures where it may precipitate a significant proportion of its sparingly soluble load. At epizonal depths, volatile saturation may occur early relative to crystallization progress because of low-prevailing load pressure. MVP saturation occurs when the sum of the vapor pressures of the components of a liquid is equal to or exceeds the load pressure;

Some Ore-metal Anomalies Formed in the Continental Realm boiling can result from a decrease in the load pressure or an increase in the vapor pressure of the magma. A decrease in load pressure for a given vapor pressure of the dissolved constituents (at constant temperature and melt composition) results in first boiling. As many melts are known to possess concentrations of water on the order of 3 – 6 wt.%, some with CO2 concentrations of up to a few thousand ppm (Lowenstern, 1994), partial pressures of magmatic gases can sum to a few hundred MPa. Volatile exsolution can also be brought about isobarically. If a liquid mixture of volatile and relatively nonvolatile components is cooled below its liquidus and begins to crystallize phases that are dominantly anhydrous, the mole fraction of the volatile components in the mixture, and, concomitantly, the vapor pressure of the magma will increase. This second mode of boiling, which is accompanied by cooling, and crystallization of nonvolatile components is termed second boiling. Some workers have suggested that mafic magma input into shallow felsic magma chambers is an important part of the volatile saturation paradigm for porphyry-type systems (Hattori and Keith, 2002). Indeed, there is ample evidence that mafic and felsic magmas interact in complex manners in subvolcanic environments (Hattori and Keith, 2002). Further, this process may be important in the budgets for chlorine and sulfur (and possibly CO2) for the ore – magma – volatile system, and significant amounts of sulfur and chlorine can be transferred from subcrustal melts into granite magma chambers, as suggested for chlorine by Piccoli and Candela (1994) on the basis of apatite chemistry in zoned intrusions in arc systems. However, it seems unlikely that the mafic magma is critical for saturation of the system with respect to a water-rich volatile phase. In summary, we might expect magma, after some differentiation at depth, to experience volatile exsolution and further crystallization upon ascent. During the phase changes occurring at each stage, ore metals and other magmatic constituents will partition among the melt, mineral, and volatile phases (whether in more primitive mafic magmas or more evolved felsic magmas), in accord with the appropriate partitioning and solubility equilibria. Therefore, the timing of volatile saturation relative to crystallization of any of the causative magma batches is critical in our understanding of ore genesis according to the orthomagmatic – hydrothermal model. The initial water concentration of a melt can only be defined for a given initial melt composition. Note also that the concentration of water in the melt phase is different from the water concentration in the magma. A magma with 2 wt.% water, 50% crystals, and which is volatile-phase undersaturated, contains a melt phase with 4 wt.% water.

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The competition for ore metals (and other elements) between the magmatic volatile phase (MVP) and the crystallizing magmatic mineral assemblage during crystallization is a first order control on the efficiency with which metals can be removed from magma into an evolving MVP or “ore fluid” (Candela and Holland, 1986). The values of the pertinent crystal/melt and MVP/melt partition coefficients, together with the timing of crystallization relative to saturation with respect to a water-rich MVP, is a major factor in determining whether the MVP or the crystallizing phases take up a higher proportion of a given ore metal. Detailed accounts of ore –metal partitioning in felsic melt –crystal systems have been published (see Candela, 1992; Jugo et al., 1999), and some conclusions are summarized here. Molybdenum and tungsten can be partitioned into titanium-bearing accessory phases such as ilmenite, magnetite (and other spinels), sphene, and biotite. Both gold and copper are partitioned strongly into some accessory magmatic-sulfide minerals (which are stable at fO2 , NNO þ 1; Carroll and Rutherford, 1985) and their protracted crystallization can result in the strong depletion of melts with respect to copper and/or gold. If a significant amount of copper- and gold-bearing sulfide is removed from the magmatic system before saturation of the melt with respect to a volatile phase, the probability of formation of a copper or gold-rich ore or protore will be significantly reduced. In granitic magmas with oxygen fugacities between QFM and NNO, a small amount of pyrrhotite crystallization might remove a significant amount of copper, but not of gold. For some metals such as copper and gold that possess high-MVP/melt partition coefficients (see Candela et al., 1996), the role of fO2 is obviated in cases where hot, primitive melts with high initial water concentrations are emplaced in the epizone. Under these conditions, volatile exsolution may occur before significant fractionation of sulfides, regardless of the oxygen fugacity. However, copper mineralization is usually restricted to relatively oxidized magmas at shallow levels, as even small amounts of pyrrhotite can remove significant amounts of copper from a crystallizing melt (Jugo et al., 1999). Candela and Holland (1986) suggested that magmatic sulfides might, in some cases, react out of the hypersolidus mineral assemblage upon saturation of a magma with a volatile phase, as sulfur is removed from the melt along with other exsolving volatiles. However, Keith et al. (1997) argue that metals sequestered in sulfides can be repartitioned into the MVP after volatile-phase saturation, if magmatic sulfides are crystallized at, or brought to, the level of emplacement where volatile exsolution occurs. This process depends on the proportion of ore metals in a magmatic

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system that are partitioned into these phases during in situ crystallization at the level of emplacement versus the proportion of ore metals that are partitioned into crystallizing phases at depths below the level of emplacement and remain at sites distal to MVP exsolution. In the latter case, ore metals are certainly lost from the system. Metals (excluding REEs) that tend to increase in concentration upon magmatic fractionation include lithium, molybdenum, tungsten, tin, rubidium, beryllium, caesium, bismuth, niobium, tantalum, manganese, uranium, and thorium. These metals therefore dominate ores associated with highly fractionated igneous rocks. However, contrary to conventional wisdom, igneous fractionation is a relatively inefficient means of forming an ore deposit. Consider a highly incompatible element, e.g., rubidium, with an initial melt concentration of 50 ppm and a bulk crystal/melt partition coefficient ,0.2. A highly evolved fractionate resulting from 90% crystallization would be considered “geochemically enriched,” with a rubidium concentration of 315 ppm. However, only 63% of the rubidium originally present would remain in the melt phase. That is, fully one-third of the rubidium would have been lost to fractionated crystallized phases. At a value for the bulk partition coefficient for rubidium equal to a more realistic 0.6 (e.g., for a melt crystallizing one-third plagioclase, onethird potassium feldspar, and one-third quartz; Cavazzini, 2001), the rubidium concentration still increases by greater than a factor of 2.5, yet fully three-fourth of the rubidium originally present in the melt would have been lost to the products of crystallization. That is, ore metals are continually dispersed throughout consolidating plutons, and hence the crust, by the action of crystallization. However, some ore metals sequestered and dispersed less readily than others, and the composition of magmatic– hydrothermal ores is dominated by those metals that suffer least in the battle of attrition waged by magmatic crystallization. In summary, early saturation with a water-rich volatile phase (Candela and Holland, 1986) allows the evolving MVP maximum access to crystal-, or sulfide liquid-compatible ore metals, such as copper or gold. In a relatively dry magma that is crystallizing at a relatively deep level, exsolution of a water-rich MVP will occur late in magmatic crystallization progress. Thus, there are two main factors affecting the efficiency of ore-metal removal from magmas: the sequestering of ore metals into crystallizing magmatic phases, and the timing of volatile saturation relative to crystallization (Candela and Holland, 1986), with early volatile exsolution allowing ore metals, especially those that are crystal compatible, to be available for partitioning into the MVP.

3.12.4.1 Epithermal Deposits Epithermal hydrothermal ore deposits form at shallow levels in volcano-plutonic arcs systems, and fall into two classes: the adularia-sericite type (also called low-sulfidation deposits) and the high-sulfidation state acid-sulfate type. These are sometimes referred to as quartz-adularia and quartz-alunite type deposits, respectively. The acid-sulfate deposits include the Nansatsu district, Japan; Goldfield, Nevada; El Indio, Chile; and Lepanto-Far Southeast, the Philippines. The low-sulfidation deposits include Creede, Colorado; McLauglin, California; Comstock and Round Mountain, Nevada; and Hishikari, Japan. Both types form in near-surface volcanic environments, and may have porphyry-type ore systems at depth. They are commonly found at shallow levels in volcanic environments including volcanic islands, continental magmatic arcs, and extensional regimes (Hedenquist and Lowenstern, 1994). The acid-sulfate deposits comprise shallowly formed veins, stockworks, breccias, or massive replacements associated with vuggy quartz and acid-leached advanced argillic-type alteration. Common metals in the anomalies include gold and copper, with arsenic, silver, zinc, lead, antimony, molybdenum, bismuth, tellurium, tin, boron, and mercury. The ores are commonly rich in pyrite and copper-bearing sulfosalts. The adularia-sericite deposits comprise shallowly formed quartz veins, stockworks, and breccias, with open-space filling textures. Common metals present in the anomalies include gold – silver – arsenic – antimony – mercury – lead – zinc – copper. Barium, fluorine, manganese, molybdenum, and selenium may also be present. Some quartz-adularia deposits have gold tellurides, and the vanadium-bearing mica, roscoelite. Heald et al. (1987) point out that the base metal contents of the acid-sulfate-type deposit are relatively high and relatively copper rich, whereas there is great variability in the base-metal contents of the adularia-sericite-type deposits (which tend to be relatively copper poor). Adularia-sericite-type gold – silver deposits are also considerably more abundant than acidsulfate-type deposits. Heald et al. further suggest that the two types of deposits form under similar conditions of depth (1 –2 km) and temperature (100 –3008) but in different parts of paleogeothermal systems. Acid-sulfate deposits form in the root zones of volcanic structures from acid waters that contain residual magmatic volatiles, whereas the adularia-sericite variants are deposited from neutral to weakly acidic, alkali chloride waters produced by the mixing of surficial waters with deeper, heated saline waters in a lateral flow regime, above and probably offset from the magmatic hearth. Boiling and mixing are two of

Epilogue the more commonly proposed mechanisms for ore deposition in the quartz-adularia deposits. The longer path length for the flow of magmatic fluids promotes both dilution of the fluid by meteoric waters, and equilibration with wall rock, both of which promote a reduction in the sulfur concentration of the ore fluid and a lowering of the oxidation state of the system. In the acid-sulfate systems, cooling magmatic vapors are titrated directly into the overlying hydrological system. Deposition occurs in acid-sulfide systems due to either cooling or mixing (dilution).

3.12.5 (GEOCHEMICALLY SIGNIFICANT) PLACERS/RESIDUAL ORES AND CRUSTAL ANOMALIES (ALUMINUM, NICKEL, GOLD) Lateritic weathering, promoted by warm-humid environments and low rates of erosion, can enhance some geochemical anomalies to the point where they may be mined at a profit. For example, lateritic weathering of dunites and peridotites or their serpentinized equivalents can produce ores of nickel with 1 –2% nickel. Nickel is commonly hosted by either iron oxides or silicates (garnierite). The deposits may be crudely vertically zoned with pisolitic iron-oxide and nickel-bearing zones above richer saprolitic silicate ore. Aluminum, the third most abundant element in the Earth’s crust, is mined as bauxite, a mixure of gibbsite, boehmite, and diaspore. These ores occur dominantly in Cenozoic deposits formed by lateritic weathering of aluminous parent rock, which has broken down to leave a high proportion of aluminum-bearing minerals. These deposits are residual in nature, with the balance of the rock constituents removed by selective leaching. These deposits form most commonly in tropical and subtropical areas. Another surficial deposit type includes the manganese-rich crusts and nodules that occur on basalt near mid-ocean ridges, and near the sediment – water interface. Whereas the ultimate source of manganese and iron (as well as the associated nickel, copper, and cobalt) is the hydrothermal alteration of MOR basalt, the fields of manganese-rich encrustations and nodules that cover large areas of sediment-starved ocean floor are the result of surficial authigenic upward remobilization of metals and the fixing of those metals at the sediment – water interface. The final case of surficial ore genesis dealt with here involves the transport of physically disaggregated residuum from physically and chemically weathered materials by wind and water, resulting in placers, mineral deposits formed at the Earth’s surface by mechanical

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concentration of dense minerals during clastic sedimentation. These processes can act as part of the normal river, lacustrine, beach, aeolian, or glacial processes that today form modern sediments, or have acted in the past to produce ancient sediments, sedimentary rocks, or their metamorphic equivalents. For example, moving water can sort particles by density, shape, etc. Common placer minerals include gold and PGE, cassiterite, garnet, corundum, diamond, wolframite, zircon, as well as magnetite, ilmenite, rutile, and other minerals. The minerals may be concentrated by hydraulic processes in marine (beach ridges or offshore bars), glacial (moraine and outwash), or in alluvial fan, fluvial or lacustrine environments. The super giant Witwatersrand detrital golduraninite ore fields formed during the Archean, when uraninite could be transported without suffering oxidation before deposition. These deposits are essentially paleofluvial placers, with ore (characterized by gold concentrations on the order of 10 ppm) occurring in extremely thin sheets with breadth to thickness ratios approaching 105 (Gilbert and Park, 1986).

3.12.6 EPILOGUE A confluence of factors determines the general distribution of the scarce metals in the Earth’s crust. First, there is the similarity of the element to a major rock-forming mineral. Possibly no element exemplifies this better than rubidium, which does not form any known mineral of its own, is commonly camouflaged in potassium minerals, and consequently is dispersed in the feldspar minerals that characterize the Earth’s crust. Even the richest rubidium ores do not contain rubidium minerals, sensu stricto; oregrade rubidium occurs in lepidolite and pollucite (potassium – lithium and caesium minerals, respectively), which occur as minor by-products in the Tanco rare-metal pegmatite deposit of Bernic Lake, Manitoba, Canada (Cerny et al., 1996). The driving force for camouflage, of course, is entropy of mixing on crystalline sites, which, at equilibrium, lowers the overall free energy of the system as long as bonding and size considerations do not preclude it. To a lesser extent, the same fate is suffered by an element such as vanadium, which is camouflaged by ferric iron, and is therefore commonly dispersed though the ferromagnesian constituents of the crust. However, vanadium does form its own minerals, and in fact does form its own deposits, even though it is commonly considered a by-product of uranium mining. This, however, may be seen as something of an economic bias, as the post-World War II industrial boom coincided with the dawn of the “nuclear age,” and the importance of uranium.

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Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 411–431

3.13 Geochemistry of the Igneous Oceanic Crust E. M. Klein Duke University, Durham, NC, USA NOMENCLATURE 3.13.1 INTRODUCTION 3.13.2 ARCHITECTURE OF THE OCEANIC CRUST 3.13.3 MANTLE MELTING: SIMPLE PASSIVE MODEL 3.13.4 WORLDWIDE GEOCHEMICAL VARIATIONS AMONG OCEAN RIDGE BASALTS 3.13.4.1 Crystallization 3.13.4.1.1 Major elements 3.13.4.1.2 Trace elements 3.13.4.1.3 Correcting for crystallization 3.13.4.2 Melting 3.13.4.2.1 Major elements 3.13.4.2.2 Trace elements 3.13.4.2.3 The shape of the melting regime and the generation of diverse melt compositions 3.13.4.3 Mantle Heterogeneity 3.13.4.3.1 Pyroxenite melting 3.13.4.3.2 Assimilation of altered crust 3.13.4.3.3 Local trends in basalt composition 3.13.4.4 Spatial Variations in Lava Compositions 3.13.4.4.1 Along-axis chemical variations 3.13.4.4.2 Temporal variations on lava composition 3.13.5 THE LOWER OCEANIC CRUST 3.13.5.1 Geochemical Systematics of Dikes 3.13.5.2 Geochemical and Textural Systematics of Gabbros 3.13.6 CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

Kd

distribution coefficient for an element in a mineral Kdij distribution coefficient for element i in mineral j Mg# 100 £ molecular [Mg/(Fe þ Mg)] Na8.0, Fe8.0, etc. Na2O, FeO, etc fractionation corrected to 8 wt.% MgO (as discussed in text) Po pressure of intersection of the solidus Pf final pressure of melting Xj fraction of the mineral j

NOMENCLATURE C0 CL Cs CS D F

433 434 434 435 437 438 438 438 443 444 444 446 447 447 448 449 450 451 451 454 456 456 457 458 459 459

the initial concentration of an element in the system concentration of an element in the melt or liquid concentration of an element in a mineral (solid) concentration of the element in the total solid bulk distribution coefficient fraction of the system that is melt 433

434

Geochemistry of the Igneous Oceanic Crust

3.13.1 INTRODUCTION Approximately 60% of the Earth’s surface consists of oceanic crust (Cogley, 1984). New ocean crust is created at divergent plate boundaries called ocean ridges or spreading centers (Figure 1). Once created, the oceanic crust is transported off-axis to each side of the spreading center, accumulating sediment as it ages, and is ultimately consumed at subduction zones and returned in a modified form to the mantle (see Chapters 7.01, 2.11, 3.15, and 3.17). The oceanic crust plays a key role in the ongoing processes that modify the compositions of major earth reservoirs. As the product of mantle melting, the generation of new oceanic crust continuously changes the composition of the upper mantle from which it forms (e.g., Chapter 2.08). In addition, the crust is the primary interface of exchange between fluids of the Earth’s surface and the solid earth below. Hydrothermal circulation of seawater through the ocean crust, for example, is a major factor controlling the chemistry of seawater (see Chapters 3.15 and 6.07). Subduction of hydrothermally altered oceanic crust is believed to initiate arc volcanism, and the particular composition of the subducting crust affects the compositions of the arc magmas (see Chapter 3.18). The deep subduction of altered ocean crust is also the primary means of recycling material back to the mantle where, convectively mixed with ambient mantle, it may form both the source

region of some hotspots and dispersed chemical heterogeneities (see Chapters 3.17, 3.18, 2.03, and 2.11). Thus, an understanding of the oceanic crust is central to our elucidation of whole-earth geochemical processes. This chapter reviews the architecture of the oceanic crust, and the geochemical processes by which it is created, including the nature and origin of its major element, trace element, and radiogenic isotopic composition. More detailed perspectives on these and related topics such as Melting and Melt Percolation Models (see Chapter 3.14), Mantle Heterogeneity (see Chapter 2.03), and the Subduction of Altered Oceanic Crust (see Chapters 3.15, 3.17, 3.18, and 2.11) are presented elsewhere in these volumes.

3.13.2 ARCHITECTURE OF THE OCEANIC CRUST Our knowledge of the architecture of the oceanic crust derives from four main sources: studies of portions of the oceanic crust that have been obducted onto land (ophiolite complexes); drilling of the ocean crust; exposures of the deeper crust at fracture zones and rare tectonic windows; and geophysical studies of the seismic properties of the ocean crust. These studies have revealed that the general structure of the igneous ocean crust, from top to bottom, consists of a carapace of basaltic lava flows and pillows

Figure 1 Schematic map of the global ocean ridge system.

Mantle Melting: Simple Passive Model (,1 –2 km thick), underlain by a sheeted dike complex (,1– 2 km thick), a gabbroic unit of diverse textures and lithologies (,3– 5 km thick), and finally ultramafic rocks of the lower crust and mantle, including harzburgite, lherzolite, and dunite (Figure 2; e.g., Penrose Conference Participants, 1972; C. Z. Mutter and J. C. Mutter, 1993; Karson, 2002). A simple view of the creation of each of these units draws upon the results of geophysical studies of melt reservoirs beneath fast-spreading ridges (Figure 3; Detrick, 1987; Sinton and Detrick, 1992) in which a shallow (,1 –2 km depth), thin (50 –100 m) relatively narrow (0.5 – 1.5 km) melt lens overlies a wider (. 8 km) mush zone (melt þ crystals) and partially solid transition zone that extends to the base of the crust (average crustal thickness is ,5 – 7 km). Melt from the mantle is focussed toward the ridge (e.g., Spiegelman and McKenzie, 1987) and replenishes both the mush zone and the melt lens, where cooling and crystallization take place. Periodically, extensional forces associated with seafloor spreading

Figure 2 (a) Generalized internal structure and interpretation of the oceanic crust derived from studies of ophiolite complexes and interpretations of marine seismic and geologic data. (b) Outcrop photographs of crustal rocks from ophiolites; top: pillow lavas, Macquarie Island; middle: sheeted dike complex, Oman; bottom: gabbroic rocks, Bay of Islands (Karson, 2002) (reproduced by permission of Annual Reviews from Annual Reviews of Earth and Planetary Sciences, 2002, 30, 347– 384).

435

create a pathway favorable for dike injection above the melt lens and, if the dike pierces the surface, magma erupts on the ocean floor as lava. Continued cooling of the melt lens and the mush zone below leads to the crystallization of the plutonic rocks of the gabbroic sequence. In addition, crystallization of mafic minerals such as olivine and pyroxene at the base of the mush zone may create a basal crustal ultramafic unit. Below the crust lie mantle rocks that often show mineralogic and chemical evidence that they are the residues of a previous melting event that provided the melt that formed the ocean crust above (Dick et al., 1984; Michael and Bonatti, 1985; Johnson and Dick, 1992).

3.13.3 MANTLE MELTING: SIMPLE PASSIVE MODEL In recent years there has been a great deal of experimental and theoretical work on the processes by which the mantle melts to form basaltic magma erupted at mid-ocean ridges (see Chapter 3.14 and Langmuir et al., 1992 for expanded reviews). In the simple passive model for mantle melting (Figure 4; e.g., Oxburgh, 1965, 1980; Lachenbruch, 1976; McKenzie and Bickle, 1988; Plank and Langmuir, 1992), viscous drag associated with seafloor spreading creates a “void” beneath the ridge axis that draws mantle up from depth (in reality, the process is viewed as two continuous conveyor belts, so no actual void space exists). If mantle melting is adiabatic, the mantle largely retains its higher temperature from depth but decompresses as it rises. In pressure – temperature space (Figure 5), this means that the mantle will intersect the solidus at some depth, depending upon its initial subsolidus temperature, and begin to melt; the slope of the melting path differs from that of the solid adiabat, because there is a temperature decrease associated with the heat of fusion. Continued corner flow causes the mantle to rise further, melting more as it ascends; thus, the amount of melting that a parcel of mantle will experience is governed by the difference in pressure between the depth of intersection of the solidus (Po) and the depth at which it turns the corner and no longer decompresses (Pf, depth of final melting). As melting proceeds, the melt separates from the solid at melt percentages less than 1% (e.g., McKenzie, 1984; Ribe, 1985; Daines and Richter, 1988; Johnson et al., 1990; Johnson and Dick, 1992; Faul, 1997), and is focused toward the ridge by processes that remain a subject of debate (Spiegelman and McKenzie, 1987; Phipps Morgan, 1987; Sparks and Parmentier, 1991). Much work has been devoted to examining diverse aspects of the assumptions in this model,

436

Geochemistry of the Igneous Oceanic Crust

Figure 3 Upper panels: Interpretive model of a magma chamber along a fast-spreading ridge like the EPR based on recent geophysical and petrological constraints. The essential elements of this model are a narrow, sill-like body surrounded by a transition zone to the solidified, but still hot, surrounding rock. The solidus, which defines the limit of magma, can occur anywhere from the boundary of the mush zone to the edges of the axial low-velocity zone (LVZ). Because the solidus may not be isothermal and significant lithologic variations can occur in the lower layered gabbros, isolated pockets of magma with low melt percentages can occur throughout the LVZ. Eruptions will mainly tap the molten, low-viscosity melt lens. The relative volumes of melt must vary along the ridge axis, particularly near ridge axis discontinuities. Lower panel: Intrepretive model of a magma chamber beneath a slow spreading (low magma supply) ridge like the MAR, based on recent geophysical and petrological constraints. Such ridges are unlikely to be underlain by an eruptable magma lens in any steady-state sense. A dike-like mush zone is envisioned beneath the rift valley forming small sill-like intrusive bodies which progressively crystallize to form oceanic crust. Eruptions will be closely coupled in time to injection events of new magma from the mantle. Faults bordering the rift valley may root in the brittle– ductile transition within the partially molten magma chamber (Sinton and Detrick, 1992) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1992, 97, 197 –216).

and their implications for the compositions of the melts produced. But, for the purposes of this review, we can use this simple passive model to explore some of the basic aspects of mantle

melting and how variables associated with melting (as well as those involved in crystallization and source composition) will affect the compositions of the melts produced.

Worldwide Geochemical Variations among Ocean Ridge Basalts

437

Figure 4 Idealized steady-state melting regimes produced by spreading of the plates. Solid curves with arrows are mantle flow paths through the melting regime. Dashed lines are contours of the extent of melting in the melting regime or the extent of melt removed in the residual mantle column (RMC). The two melting regimes are for two mantle temperatures. Hotter mantle intersects the solidus deeper (see Figure 5), leading to greater extents of melting, a taller RMC, and thicker crust (Langmuir et al., 1992) (reproduced by permission of American Geological Union from J. Geophys. Res., 1992, 71, 183 – 280).

Figure 5 Schematic pressure (P) versus temperature (T) diagram showing melting paths for adiabatically ascending mantle. Melting contours are evenly spaced for illustration only. The two paths are for two different mantle temperatures. The further a mantle parcel ascends, the greater the extent of melting (Langmuir et al., 1992) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1992, 71, 183– 280).

3.13.4 WORLDWIDE GEOCHEMICAL VARIATIONS AMONG OCEAN RIDGE BASALTS The vast majority of lavas erupted along the mid-ocean ridge system are basalts. This uniformity of rock type, however, belies a great diversity in composition resulting from the disparate physical and chemical conditions by which, and from which, the magmas are produced. A number of compilations of worldwide mid-ocean ridge basalt (MORB) compositions now exist, the most complete of which is RidgePetDB (petrologic database of the ocean floor; Lehnert et al., 2000)

available at http://petdb.ldeo.columbia.edu/petdb. Other oceanic basalt databases are available through the Smithsonian Institution basalt glass data file (www.hrw.com/science/si-science/ physical/geology/mineralsci/research/glass/); and the GEOROC (geochemistry of rocks from the oceans and continents) database at http://georoc. mpch-mainz.gwdg.de/Start.asp. Select examples of various MORB compositions are presented in Tables 1 and 2. A wide range of processes produces the diversity of magma compositions erupted along mid-ocean ridges. Principal among these are variations in source composition, in the style of mantle melting, and in the crystallization of the magma en route to the surface. These processes produce identifiable chemical signatures in terms of major element, trace element, and radiogenic isotope compositions, which are discussed in the following paragraphs. A common approach to studying the chemical variations in MORB begins by identifying the effects of the last magmatic processes to occur, namely, crystallization. Chemical systematics that cannot be accounted for by crystallization are then examined from the perspective of melting. Finally, those aspects of chemical variability that cannot be attributed to either crystallization or melting are then explored in terms of variations in source composition. To be sure, there is ample evidence for feedback among the three processes: source composition itself likely affects melting systematics, and melting variations can affect the style of crystallization (and additional processes, such as magma chamber contamination by seawater, may occur; e.g., Michael and Cornell, 1998). But it is only by initially examining each of the three main processes in isolation that the

438

Geochemistry of the Igneous Oceanic Crust

Table 1 Five basaltic glasses from the East Pacific Rise near the Clipperton Transform that, to first order, form a coherent liquid line of descent for fractional crystallization from the highest MgO magma (#1) to lowest (#5). 1 SiO2 TiO2 Al2O3 FeOT MnO MgO CaO Na2O K2O P2O5 Sum La Ce Nd Sm Yb Ba Nb Ni Sr Y Zr

50.46 1.28 15.55 8.93 0.19 8.37 12.21 2.47 0.11 0.11 99.7

50.65 1.53 15.15 9.88 7.85 11.95 2.54 0.08 0.18 99.8

49.8 1.64 14.28 10.45 0.2 7.21 11.68 2.59 0.13 0.1 98.1

3.38 10.6

3.9 12.09 11.1 3.36 3.77 3.18 3.59 7.43 12.1 3.06 80 72 120 122 38.9 100 111

Mg# 87

2 3 4 5 More primitive ! more evolved

62.6

58.6

86

Sr/ Sr Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb

55.2 0.702505 0.513139 18.312 15.484 37.722

143

50.43 2.07 14.65 11.10 0.20 6.87 11.40 2.71 0.19 0.25 99.9

50.93 2.55 13.51 13.13 0.25 5.62 10.24 3.2 0.2 0.22 99.9

5.46 5.86 16.8 18.2 14.3 16 4.64 5.41 4.14 4.75 20.4 20.7 4.89 5.13 71 41.4 134 133 39.7 43.9 135 155 52.5

43.3 0.702566 0.513155

Source: samples (left to right) WASRAI2-050-007; WASRAI2-057006; ODP0142-0864A-001M-003/0-3; MELPHNX-2-068-001; MELPHNX-2-GC083 from RidgePetDB (Lehnert et al., 2000). Major element oxides are in wt.%; trace elements are in ppm. FeOT is total iron.

interplay among them (and the influence of other factors), can then be overlaid. Thus, we begin by examining the last magmatic process, crystallization, first; it is also the process that is probably best understood and quantified. 3.13.4.1 Crystallization 3.13.4.1.1

Major elements

It has long been known that suites of MORB samples recovered from a relatively limited spatial area often display regular variations in major element oxides as a function of MgO content (Figure 6) or Mg# (100 £ molecular [Mg/(Fe þ Mg)]). In addition, the Mg#s of most MORBs (mean Mg# ¼ 55 ^ 7.6 (1s); n , 1.3 £ 104) are not in equilibrium with presumed mantle olivine compositions (Mg# , 90; Roeder and Emslie, 1970; Langmuir and Hanson, 1981). These findings, combined with the results of experimental

studies on the crystallization of basaltic melts (Bender et al., 1978; Walker et al., 1979; Stolper, 1980; Grove and Bryan, 1983), have led to the recognition that basaltic melts commonly undergo significant and varying extents of crystallization in sub-axial magma reservoirs prior to eruption. Magmas at varying stages of crystallization are periodically tapped from these reservoirs, and erupted at the surface, revealing the evolution of the melt as it cools and crystallizes. The magnitudes of the changes in each oxide as a function of decreasing MgO result from the crystallization from the melt of stable mineral phases (which, for typical basaltic melts at ,50 – 100 MPa pressure, generally includes olivine ^ minor spinel, followed by plagioclase and finally clinopyroxene). Thus, in Figure 6, the compositions and proportions of the minerals that crystallize from each melt drive the residual melt to a different (lower MgO) composition. Because these changes represent the changing composition of the melt as it cools, crystallizes, and evolves, the variations are called the “liquid line of descent” for a particular parental (high MgO) magma composition; a calculated liquid line of descent is shown in Figure 6, which closely approximates the trend of the data. A number of models now exist that calculate the liquid line of descent for crystallization of a given parental magma composition under specified conditions of pressure, oxygen fugacity, and other variables (e.g., Nielsen, 1985, 1990; Weaver and Langmuir, 1990; Longhi, 1991; various versions of the MELTS program available on the Web: www.geology.washington.edu/~ghiorso/ MeltsWWW/Melts.html, Ghiorso and Sack, 1995; Ghiorso et al., 2002). These model calculations suggest that, often, much of the major element variability among MORB from a given region can be accounted for by low (,50 MPa) to intermediate (e.g., 200– 300 MPa) pressure fractional crystallization (evidence for high-pressure crystallization, at 600 –800 MPa, has also been described (e.g., Elthon and Scarfe, 1984; Grove et al., 1990) but is believed to be less common). During fractional crystallization, the crystallizing minerals separate from the cooling magma as they are formed; this differs from batch or equilibrium crystallization in which minerals and melt remain in contact and continue to equilibrate as cooling and crystallization proceed. Table 1 presents data for magmas that, to first order, appear to result from progressive fractional crystallization from the most primitive (highest MgO) magma in the suite. 3.13.4.1.2

Trace elements

For more than three decades, geochemists have studied the trace element variations in MORB

Table 2 Representative analyses of MORBs. Kolbeinsey Ridge, northern EPR, and mid-Cayman Rise encompass much of the global range in extents and pressures of melting. Also shown are examples of an Indian Ocean MORB and composite analyses of a normal (N-MORB), a transitional (T-type) MORB, and an enriched (E-type) MORB from the MAR. Global range: N-MORB

Indian Oc. MORB

MAR N-MORB

MAR T-MORB

MAR E-MORB

50.88 1.71 16.07 9.73

Kolbeinsey ridge N. EPR N-MORB Mid-Cayman Rise ((high F, high P) left to (low F, low P) right) SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2 O P2O5 Sum

50.55 0.82 14.83 10.49 0.18 8.53 12.99 1.50 0.05 0.06 100.0

50.16 1.47 15.79 9.51 0.16 7.58 12.19 2.76 0.13 0.13 99.9

50.98 1.72 16.11 8.745 0.17 7.23 10.34 3.56 0.24 0.22 99.3

50.85 1.10 16.72 7.76 0.15 8.85 11.11 3.17 0.15 0.13 100.0

50.01 1.11 16.31 9.73 0.14 8.67 11.75 2.52 0.05 0.08 100.4

7.39 11.17 2.89 0.11 0.16 100.1

51.28 1.83 15.23 9.60 0.16 7.43 10.59 3.08 0.53 0.26 100.0

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Ba Co Cr Cs Cu Hf

1.38 4.10 0.74 4.32 1.69 0.68 2.44 0.48 3.39 0.76 2.30 0.34 2.30 0.340 12.0 54.0 161

3.34 10.40 1.91 9.62 3.14 1.18 3.97 0.72 4.85 1.03 2.72 0.38 2.63 0.400 12.2 50.1 253 0.027 82.3 2.14

5.77 17.10

4.46 12.21

1.88 5.99

5.34 13

11.5 26

4.35 1.53

8.47 2.72 1.07 3.49

6.07 2.22 0.9 3.5

10.7 3.39 1.42 4.32

17.1 4.38 1.54 5.26

4.04

4.46

5.57

5.24

2.56

2.57

3.24

2.83

2.25 0.354 19.3

2.72

3.28

2.68

446 0.014 57.0

251 0.006 68 2.9

137.0 1.44

0.93

3.42 0.469 22.0 37.8 250 3.70

6.11

31

123

0.04

0.08

2.12

2.06 (continued )

Table 2 (continued). Global range: N-MORB

Indian Oc. MORB

MAR N-MORB

MAR T-MORB

MAR E-MORB

5.15 115 2.34 43 129 0.38 0.45 0.11 299 32

11.2 91 0.95 8.85 36 181 0.77 1.07 0.28 288 29

100 0.70268 0.51296 18.71 15.567 38.478 25.48 N 45.38 W 3200

134 0.70392 0.51271 18.032 15.507 38.237 47.978 S 10.088 W 2895

Kolbeinsey ridge N. EPR N-MORB Mid-Cayman Rise ((high F, high P) left to (low F, low P) right) Nb Ni Pb Rb Sc Sr Ta Th U V Y Zn Zr 87 Sr/86Sr 143 Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb Latitude Longitude Depth (m)

3.08 81 0.155 1.07 48.0 68

20.6 80.0 39 0.7029 0.51319 17.953 15.424 37.608 67.988 N 18.318 W 740

2.99 120 0.359 1.45 42.3 142 0.203 0.141 0.061 265 27.2 80.2 89 0.7026 0.51316 18.286 15.464 37.636 118 22.80 N 1038 39.60 N 3063

5.02 123

220 36.0

200 24.4

1.07 119 0.19 0.38 44 94 0.13 0.09 0.03 281 25

153 0.7025 0.51315 18.21 15.48 37.66 18.438 N 81.728 W 4795

92 0.7035 0.51299 17.764 15.483 37.803 49.918 S 115.388 E 3087

57 0.7025 0.51306 18.215 15.535 38.047 49.768 S 88 W 3874

33.6 188 0.309 0.18

137 1.08 30.3 191

Data sources are as follows: Kolbeinsey Ridge: RidgePetDB (Lehnert et al., 2000) for sample POS0158-404-00; major elements and most trace elements on whole rock powders; Pb, Sr, Rb, and isotope ratios on glasses. N. EPR: glass reported by Niu et al. (1999); mid-Cayman Rise: glass compositions reported in RidgePetDB for sample KNO0054-027-005, augmented with Ba, V, and Y data on a similar sample reported by Thompson et al. (1980) and the sole isotopic analysis of a mid-Cayman Rise basalt from RidgePetDB; Indian Ocean MORB glass reported by Klein et al. (1991); N-Type, T-type, and E-Type MAR MORB glass are composite analyses reported in the RidgePetDB database (major and most trace elements for MAR N-type, T-type, and E-type MORB are respectively from samples: EW19309-012-00, VEM0025-001-022 and EW19309-004-002; note that latitude, longitude, and depth for the composite analyses refer to the locations of the samples for which the major element data are reported). Major element data for the Kolbeinsey Ridge, Indian Ocean, and MAR T-type samples have been recalculated to express Fe2O3 as FeO. Major element oxides are in wt.%; trace elements are in ppm.

Worldwide Geochemical Variations among Ocean Ridge Basalts

441

Figure 6 MgO variation diagrams for microprobe data for basalt glasses from near the Clipperton Fracture Zone (from RidgePetDB; Lehnert et al., 2000). Also shown is a calculated liquid line of descent for fractional crystallization at ,50 MPa using the program of Weaver and Langmuir (1990) (labeled W þ L) for a primitive composition at ,9 wt.% MgO. Oxides in wt.% (Langmuir et al., 1992) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1992, 71, 183 – 280).

to understand the magmatic processes that occur in their genesis. Pioneering work by Gast (1968), Shaw (1970) and Schilling (1966, 1971) laid much of the theoretical and computational foundation for trace element modeling that continues to be used and expanded upon today. The basic elements of this approach can be summarized as follows. Because much of the chemical variability among the MORB results from the fractionation of melt and minerals, it is useful to approach the partitioning of a given trace element between solid and liquid (melt) in terms of its distribution coefficient, Kd: Kd ¼ Cs =CL

ð1Þ

where Cs is the concentration of an element in a particular mineral (solid) and CL is its concentration in the liquid (melt). Distribution coefficients for various elements between minerals and melts of different compositions are generally determined experimentally and are available for most of the common minerals involved in mantle melting and

basalt crystallization (see, e.g., http://earthref.org/ GERM/; and Chapter 2.09). For a multi-mineral system, the bulk distribution coefficient, D, is the combination of the Kd for each mineral weighted by the fraction of each mineral in the system, and is defined as follows: D ¼ ½ðfraction mineral 1Þ £ ðKd for mineral 1Þ þ ðfraction mineral 2Þ £ ðKd for mineral 2Þ þ etc: X ¼ Xj Kdij

ð2Þ D ¼ CS =CL

ð3Þ Kdij

where Xj is the fraction of the mineral j; is the distribution coefficient for element i in mineral j; and CS is the concentration of the element in the total solid. From the formulation above, it is clear that for a multi-mineral solid/melt system, an element that has a D . 1 is preferentially concentrated in

442

Geochemistry of the Igneous Oceanic Crust

the solid; such elements are described as “compatible” because they are such in the mineral assemblage present. An element that has a D , 1 is preferentially concentrated in the melt, and is described as an “incompatible” element. (And an element with a D ¼ 1 will exist in equal concentration in the solid and melt.) During any closed-system process, mass balance must be conserved: C0 ¼ ðFCL Þ þ ½ð1 2 FÞCS 

ð4Þ

where C0 is the initial concentration of an element in the system and F is the fraction of the system that is melt. By substituting (3) into (4), we obtain CL =C0 ¼ 1=½F þ Dð1 2 FÞ

ð5Þ

This equation, when applied to melting, is called the “equilibrium melting” or “batch melting” equation and, when applied to crystallization, it is called the “equilibrium crystallization” equation. As emphasized by Langmuir et al. (1992) this equation, and others like it, is enormously powerful and can be applied to model both major elements and trace elements. This approach has received its most widespread use, however, in modeling trace element variations. A common application of this equation to trace element modeling is to examine the variations in trace element abundances and ratios for elements with different bulk distribution coefficients (Figure 7). In this plot, F is the fraction of melt: for equilibrium crystallization, F proceeds from

1 (all melt) to 0 (all solid); for equilibrium (or batch) melting, F proceeds from 0 (all solid) to 1 (all melt). During crystallization, a highly compatible element (e.g., D ¼ 10) is incorporated in the crystallizing mineral(s) and depleted in the melt, such that the final liquid present (F approaching 0) will have a concentration of 1/D, less than its starting concentration. In contrast, a highly incompatible element (e.g., D ¼ 0.01) during equilibrium crystallization gradually increases in concentration, and at large extents of crystallization (e.g., F , 0.2) becomes highly concentrated, again approaching 1/D, a concentration far greater than the original starting concentration. Compare these results for equilibrium crystallization to those produced by fractional crystallization (Figure 8(a)). Note, for example, that during fractional crystallization, the concentration in the melt of a compatible element is rapidly depleted as the crystallizing assemblage removes this element from the system. In addition, after significant fractional crystallization, moderately incompatible elements (e.g., D ¼ 0.1) are enriched in the melt in excess of the amount that can be achieved through equilibrium crystallization. Thus, by examining the concentrations of compatible and incompatible elements as a function of indicators of crystallization (e.g., MgO or Mg#) the effects of crystallization can be identified as well as the type of crystallization that has occurred. Another common application of this approach is to examine the ratio of two different incompatible elements with different Ds. This involves solving Equation (5) for the ratio of two different

Figure 7 Illustration of the effects of equilibrium (batch) crystallization or melting on trace element abundances. (a) Variation in liquid concentration (CL) (normalized to unit source concentration C0 ¼ 1) as a function of melt fraction (F) for six “elements” with different bulk distribution coefficients (D). (b) Change in the ratios of incompatible elements with different Ds as a function of F. Each curve is for a different pair of “elements” that have the Ds indicated. Note that when D , 0.1, incompatible element ratios can be changed only at very low extents of melting (or high extents of crystallization) (Langmuir et al., 1992) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1992, 71, 183 –280).

Worldwide Geochemical Variations among Ocean Ridge Basalts

443

changes little until F is less than ,0.5; at very small remaining melt fractions, the ratios increase abruptly, particularly for ratios of more moderately incompatible elements. Because the ratios of incompatible elements change little during the first 50% of crystallization (most MORB experience less than 50% crystallization), a relative constancy of such ratios with decreasing MgO content is expected. Thus, large variations in this ratio accompanied by major element indices suggesting only modest extents of crystallization (e.g., moderate to high MgO or Mg#), must be attributed to other processes, such as variations in extent of melting or source composition. A number of more complex crystallization processes have been envisioned. One that has attracted particular attention, in large part because it is inherently linked to a realistic physical model for crystallization, is called in situ crystallization (Langmuir, 1989; Sparks et al., 1984; Sparks, 1989). This process recognizes that in a crystallizing magma body, there is likely to be a temperature gradient from hottest in the center of the magma body to coolest along its margins; thus, crystallization is most likely to take place within a boundary layer between magma and country rock. As crystallization proceeds within this boundary layer, more evolved magma may be expelled and mix with the main body of magma, where it is periodically tapped and erupts lavas. Because the magma expelled from the boundary layer may be highly evolved, it may produce unusual enrichments in incompatible trace elements and ratios that exceed those produced by fractional crystallization alone.

3.13.4.1.3

Figure 8 Illustration of the effects of fractional crystallization and melting on trace element abundances. Notation as in Figure 7. (a) Fractional crystallization for five different values of D. (b) Fractional melting. (c) Accumulated fractional melting. Note the very similar results produced by accumulated fractional melting to the equilibrium melting results of Figure 7 (Langmuir et al., 1992) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1992, 71, 183– 280).

elements ðCL1 =CL2 ¼ C01 =C02 ½ðF þ D2 ð1 2 FÞ= ðF þ D1 ð1 2 FÞ: The results of this calculation, shown in Figure 7(b), reveal that for equilibrium crystallization, the ratio in the melt of two elements that have similar incompatibility

Correcting for crystallization

The discussion above emphasizes that lavas recovered from a limited length of ridge often exhibit fairly regular variations in both major and trace elements as a function of decreasing MgO (or Mg#), and these are commonly attributed to the effects of low (,50 MPa) or intermediate (300 MPa) pressure fractional crystallization. These chemical variations resulting from crystallization obscure differences in parental magma compositions that exist from region to region. Thus, methods have been developed to minimize the effects of fractionation on each oxide or trace element. One method projects the composition of each magma along a presumed slope of the liquid line of descent to a constant value of MgO, in this case 8 wt.% (Klein and Langmuir, 1987; see also updated algorithms in Castillo et al., 2000). The calculated Na2O value at 8 wt.% MgO is therefore referred to as Na8.0. In theory, if the liquid line of descent is known for any major or trace element,

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a “fractionation-corrected” value of the element can be calculated (producing calculated values of, e.g., Fe8.0, Al8.0, K8.0, Ce8.0, etc.). The few elements whose liquid lines of descent change in the sign of the slope as the fractionating phases change, such as CaO (due to the appearance of calcium-bearing phases on the liquidus), are more difficult to model in this way and therefore produce less reliable fractionation-corrected values. Nevertheless, most major and trace elements display fairly regular liquid lines of descent, and their fractionation-corrected values reveal relative differences in parental magma compositions that result from differences in melting systematics or source composition (see, e.g., differences in average Na8.0 for samples from the regions shown in Figure 9).

3.13.4.2 Melting 3.13.4.2.1

Major elements

One of the startling findings revealed by examining regional averages of fractionationcorrected values of major element compositions is that the major elements do not vary independently of one another but rather, to first order, correlate in predictable ways. In general, regions with low mean Na8.0 are also characterized by high mean Fe8.0 (as well as low Si8.0 and Al8.0, and higher Ca8.0), while other regions exhibit the opposite characteristics, as well as a continuum of compositions in between (Klein and Langmuir, 1987; Langmuir et al., 1992). Furthermore, these major element variations have been shown to correlate with physical characteristics of the ridge axis from which they were recovered. Regional averages of Na8.0 and Fe8.0, for example, show a positive and an inverse correlation, respectively, with the average ridge depth from which the lavas were recovered (Figure 10). In addition, Na8.0 was also shown to correlate inversely with seismically and geologically determined estimates of the thickness of the oceanic crust in each region (Figure 11). Thus, some of the most fundamental physical and chemical parameters studied at ocean ridges suggest a common origin in their variability. The chemical systematics can be understood as the interplay of two main factors affecting the style of mantle melting: the extent of melting and the pressure of melting. Elements such as sodium are moderately incompatible during melting of the mantle minerals (D , 0.02– 0.03), and therefore will be concentrated in the melt at small extents of melting. Iron is well known to vary strongly in the melt as a function of the pressure of melting (e.g., Langmuir and Hanson, 1980). Thus, the inverse correlation between mean Na8.0 and mean Fe8.0 would suggest that there is a positive correlation between the mean extent of melting

Figure 9 (a) MgO versus Na2O in MORBs from five regions. Samples are from the Mid-Cayman Rise, from near the Kane fracture zone at 23.22 – 23.678 N on the MAR, from the AMAR region at 36.42 – 6.798 N on the MAR, from near the Clipperton fracture zone on the EPR, and from the Kolbeinsey Ridge north of Iceland (data from the Smithsonian catalogue of basalt glasses (Melson et al., 1977), except for Clipperton data (RidgePetDB; Lehnert et al., 2000), and Kolbeinsey data (Schilling et al., 1983)). Lines show slope used by Klein and Langmuir (1987) to calculate Na8.0 (Na2O at 8% MgO). (b) Same data with calculated LLDs for a primitive composition from each of the five suites using the program of Weaver and Langmuir (1990) for fractional crystallization at 50 MPa. Vertical bar indicates the point at which clinopyroxene (cpx) joins olivine and plagioclase (ol – pl) as a fractionating phase, at which point the slope of the LLD steepens (after Langmuir et al., 1992). Original algorithms for calculation of Na8.0 and Fe8.0 for samples with 5.0– 8.5 wt.% MgO are as follows (Klein and Langmuir, 1987): Na8.0 ¼ Na20.373 £ (MgO) 2 2.98; Fe8.0 ¼ FeO þ 1.664 £ (MgO) 2 13.313. More recently presented algorithms (that permit calculation for samples with MgO . 5.0 wt.% and that take into account the changing slope of the liquid line of descent are as follows (from Castillo et al., 2000): Na8.0 ¼ 0.6074 2 3.523 £ [(Na2O þ 0.00529 £ MgO2) 2 0.9495)/(MgO 2 0.05297 £ (MgO2)28.133)]; Fe 8.0 ¼ 1.8252 1.529£ (FeO 2 0.03261 £ (MgO2) þ 0.2619)/ (MgO 2 0.04467 £ (MgO2) 2 6.67)] (Langmuir et al., 1992) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1992, 71, 183 – 280).

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Figure 11 Regional averages of seismically determined crustal thickness versus Na8.0 (the Na2O content of basalts normalized to 8 wt.% MgO; Klein and Langmuir, 1987). Sources for seismic determinations of crustal thickness are from Klein and Langmuir (1987), augmented and/or superceded by the following: Smallwood and White (1998); Navin et al. (1998); Darbyshire et al. (2000); Detrick et al. (2002); Muller et al. (1999); Hooft et al. (2000); Fowler and Keen (1979); Canales et al. (1998); McClain and Lewis (1982); Kodaira et al. (1997); Klingelho¨fer et al. (2000); Jokat et al. (2003); Michael et al. (2003). Chemical analyses are primarily on basalt glasses from the Ridge Petrologic Database (Lehnert et al., 2000) (reproduced by permission of Nature Publishing Group from Nature, 2003, 423, 932 – 933). Figure 10 Global correlations between regional averages of smoothed axial depth and the Na8.0 and Fe8.0 of MORBs. This figure is the same as figure 2 in Klein and Langmuir (1987), except all analyses have been recalculated to sum to 100% (all iron as FeO) and have been corrected for interlaboratory differences. Corrections were applied to all analyses determined at the Smithsonian Institution (Melson et al., 1977) or that were reported to be consistent with the Smithsonian data (i.e., Schilling et al., 1985), based on the factors reported in Klein and Langmuir (1989; their figure 5), so that all data are consistent with those determined at Lamont and by wet chemistry at URI. Data sources as in Klein and Langmuir (1987), with updates from the Smithsonian Catalogue of Basalt Glasses and RidgePetDB (Lehnert et al., 2000). Different groups of MORB are distinguished: solid boxes are MORB from “normal” ridge segments; diamonds are from back-arc basins; open squares are from ridges influenced by the Galapagos, Azores, Jan Mayen, Tristan, Iceland, and Bouvet hotspots; Xs are from ridge segments immediately adjacent to these hotspots. Note that Iceland and adjacent ridges are also coincident with the normal MORB array (Langmuir et al., 1992) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1992, 71, 183– 280).

and the mean pressure of melting. Melts from bathymetrically deep ridges tend to be produced by smaller extents of melting (high Na8.0) and low pressure (low Fe8.0), while melting beneath

shallower ridges leads to larger extents of melting (low Na8.0) at higher pressures of melting (high Fe8.0). It is important to recognize that the parameter “the extent of melting” has real physical implications for the volume of melt produced. Assuming mantle source regions of approximately equal volumes, and that most of the melt produced segregates to form the crust, a region that experiences a small extent of melting would be expected to have thinner oceanic crust than a region that experiences a larger extent of melting. This, in turn, would lead to the observed correlation between a chemical parameter indicative of the extent of melting (e.g., Na8.0) and a physical manifestation of the amount of melt, crustal thickness (Figure 11). Furthermore, if the crust is isostatically compensated, thinner crust would lead to greater ridge depth below sea level, and therefore the correlation between chemistry and axial depth (Figure 10). The “global correlations” among regional averages of major elements and between these elements and physical parameters of the ridge are likely due to lateral variations in the sub-solidus temperature of the mantle (Klein and Langmuir, 1987; McKenzie and Bickle, 1988). Regions underlain by hotter sub-solidus mantle intersect the solidus deeper and melt more upon ascent than

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regions underlain by cooler sub-solidus mantle temperatures (Figures 3 and 4), leading to the associated correlations with axial depth and crustal thickness. While a portion of the chemical variability among MORB may result from chemical and/or mineralogical heterogeneity of the mantle (e.g., Shen and Forsyth, 1995), the weight of current evidence supports the idea that the firstorder trends of the global correlations of regional topography, basalt chemistry, and crustal thickness result from mantle temperature variations. The discussion of melting above has focused on “mean” values (e.g., mean pressures of melting). It is likely, however, that melting in a given region occurs over a range of pressures, from the solidus to the lithosphere, and that small melt fractions segregate from their source, pool as they rise, and accumulate in shallow crustal magma reservoirs (McKenzie, 1984, 1985; Klein and Langmuir, 1987; McKenzie and Bickle, 1988; Plank and Langmuir, 1992). Thus, melting must be viewed as polybaric; the accumulated melt is not in equilibrium with mantle at any one pressure, but rather represents a mixture of melts derived from various pressures. Accordingly, accumulated melts that exhibit a higher mean pressure of melting than those from another region acquire their higher-pressure signatures by beginning melting at greater depths, hence a deeper solidus. The depth of the solidus also governs the extent of melting; a deeper solidus means that mantle will decompress and melt over a larger range of pressures as it rises (Figures 3 and 4). Accumulated melts from this deeper melting “column” will exhibit chemical characteristics indicative of larger mean extents of melting (e.g., low Na8.0) and higher mean pressures of melting (higher Fe8.0), as is observed in the global correlation of regional averages (Figure 10). It is also important to note that the eruption on the surface of melts generated at high pressures (e.g., high Fe8.0) suggests that these melts rise from great depths without experiencing significant re-equilibration en route to the surface. One explanation for the apparent absence of reequilibration may be that the melts do not rise by diffuse porous flow but rather rapidly segregate into semi-isolated channels. Indeed, Kelemen et al. (1995) have interpreted dunite channels in the mantle section of the Oman ophiolite as the residual conduits through which chemically isolated melt flow occurred. A number of studies have used the major element (and trace element) systematics of MORB to constrain the mean extent and mean pressure of melting, as well as the depth and temperature of intersection of the mantle solidus, both for a given region and for the global ocean ridge system as a whole (e.g., Klein and Langmuir, 1987;

McKenzie and Bickle, 1988; Langmuir et al., 1992; Kinzler and Grove, 1992; Asimow et al., 2001; Asimow and Langmuir, 2003). A thorough evaluation of these parameters requires numerous assumptions and information that are subject to uncertainty. These include, for example, assumptions about the physical form of the melting regime including variables such as active versus passive upwelling and variations in the final depth of melting (e.g., Scott and Stevenson, 1989; Plank and Langmuir, 1992); the processes of melt extraction and mixing (e.g., batch versus fractional melting, percentages of melt retention, incomplete focusing of melt); a melt generation function (variously calculated as ,2.5– 20% melt per GPa pressure release as a function of mineralogy and melt composition; e.g., Ahern and Turcotte, 1979; Cawthorn, 1975; McKenzie, 1984; Hirschmann et al., 1999; Asimow et al., 1997), and the effects of source mineralogy and composition including volatile species (e.g., Langmuir et al., 1992; Asimow et al., 2001; Asimow and Langmuir, 2003). With these caveats in mind, studies noted above have estimated the global range (from cold to hot regions) in the pressure of intersection of the solidus as ,1.5 – 3.5 GPa, in the temperature of intersection of the solidus as ,1,300 –1,550 8C, in the mean extent of melting as ,8 – 22%, and in the mean pressure of melting as ,0.5– 1.6 GPa. 3.13.4.2.2

Trace elements

Moderately incompatible and compatible trace elements also support the conclusions regarding extents and pressures of melting noted above. Regional averages of fractionation-corrected values of cerium (a moderately incompatible element), for example, correlate positively with Na8.0 (Figure 12(a)) consistent with regional variations in the extent of melting. In contrast, scandium (an element that is compatible in clinopyroxene in the mantle; Figure 12(b)) shows an inverse correlation with Na8.0, reflecting the fact that at small extents of melting (high Na8.0) residual clinopyroxene in mantle lherzolite retains much of the scandium; as melting proceeds and clinopyroxene melts out of the residue, scandium increases. (Note that little correlation is found between Na8.0 and Ba8.0 (Figure 12(c)), reflecting the variations in source composition of the highly incompatible element barium). Trace element studies of abyssal peridotites, the residues of mantle melting, suggest further that melting must be nearfractional (“incremental” with porosities ,1%), in order to produce the observed fractionation between elements with similar distribution coefficients such as Ti/Zr or Lu/Hf (Johnson et al., 1990; Johnson and Dick, 1992).

Worldwide Geochemical Variations among Ocean Ridge Basalts

Figure 12 Na8.0 versus (a) scandium (in ppm), (b) Ce8.0, and (c) Ba8.0 for regionally averaged data. Each point is the average of samples from an individual region, and the error bars are one standard deviation of the values. Ce8.0 and Ba8.0 (considered as highly incompatible elements) are calculated using log(Ba8.0 or Ce8.0) ¼ log(Ba or Ce) 2 0.11 £ (8 2 MgO). Regions are: Kolbeinsey Ridge (Kb), 178 N in the Red Sea (Red), near Tamayo transform of the EPR (Tam), Juan Fernandez microplate (JF), near Kane of the MAR (Kn), 13 – 148 N on EPR (13 – 14), 8 – 98 N on EPR (8 –9), south Explorer Ridge (EXP), Zones A, B, and C of the Australian Antarctic Discordance (ZA, ZB, ZC), near the Rodriguez or Indian Ocean Triple Junction (RTJ), and around 308 S on the Southwest Indian Ridge (SWIR) (Klein and Langmuir, 1987) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1987, 92, 8089– 8115).

3.13.4.2.3

The shape of the melting regime and the generation of diverse melt compositions

A paradigm has emerged that combines a view of the physical process of mantle upwelling and melting with the pooling of melts as they are produced. In a simple cornerflow model (Ahern and Turcotte, 1979; Phipps Morgan, 1987; McKenzie and Bickle, 1988; Plank and Langmuir, 1992), melts are produced throughout a wide triangular-shaped melting regime wherever mantle rises (Figure 4). A diversity of melt compositions

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are produced, resulting from the different pressures, extents and styles of melting, and the changing source composition at each point within the melting regime. Because the residual porosity is believed to be low (,1%), these melts rapidly segregate from the solid mantle and pool and mix into channels as they are focused toward the ridge axis, where they then accumulate and further mix in crustal magma reservoirs. Thus, a diversity of melt compositions are believed to be produced in the mantle, but the melt erupted at the ridge is the pooled product of these diverse melts. Evidence of the diversity of melt compositions that are generated in the melting regime comes from two main sources. The first is the study of the geochemistry of near-axis seamounts, which are believed to tap the less-pooled products of melting beneath the adjacent ridge (e.g., Allan et al., 1987; Zindler et al., 1984; Graham et al., 1988; Batiza and Vanko, 1984; Batiza et al., 1996; Niu and Batiza, 1997; Karsten et al., 1990; Hekinian et al., 1989; Reynolds et al., 1992). These studies show that lavas erupted along near-axis seamounts record a wider range of magma compositions than are erupted along the adjacent ridge, including both more depleted and more enriched compositions (see, e.g., Figure 13). The second line of evidence comes from studies of melt inclusions. Melt inclusions are small pockets of melt that are trapped within phenocrysts as the minerals crystallized. Thus, to the extent that the phenocrysts crystallized from a less evolved or less mixed melt, the melt inclusions record the nature of more primitive melt compositions. Studies of melt inclusions have revealed that although they are broadly similar in major element composition to their host lavas, they tend to extend to more primitive (higher MgO or Mg#) compositions, suggesting that indeed they record an earlier point in the evolutionary history of the magma (e.g., Sobolev and Shimizu, 1993; Sobolev, 1996; Sinton et al., 1993; Nielsen et al., 1995; Shimizu, 1998; Sours-Page et al., 1999, 2002). Notably, however, there is often a decoupling of trace element compositions from major element compositions, such that melt inclusions as a whole span a greater range in incompatible element abundances and ratios (Figures 14 and 15). Thus, melt inclusions attest to the greater diversity of melt compositions that exist in the mantle prior to pooling and eruption of magma at the ridge.

3.13.4.3 Mantle Heterogeneity The discussion above presupposes that mantle heterogeneity does not play such a major role that it obscures or dominates the major element signatures of different extents and pressures of

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Figure 13 TiO2 (wt.%) versus K2O/TiO2. Large symbols are for samples from the EPR between 118 450 N and 158 N. Note that the highest MgO lavas in each sample/symbol group have lowest TiO2; within each sample group, TiO2 increases with decreasing MgO (increasing extents of crystallization); seamount lavas are shown by crosses. Also shown are model curves for 3 – 30% batch partial melting of peridotite (lower dashed line) and of two hypothetical enriched pyroxenite compositions (upper dashed curves). Two mixing lines (solid lines) are shown extending from a peridotite melt to the two pyroxenite melting curves; these mixing lines are parallel to a line formed by the most primitive (low TiO2) lavas within each group, supporting the view that the most primitive lavas from each group may form by variable degrees of mixing between peridotite and pyroxenite melts. However, uncertainty in the composition of the enriched pyroxenite bulk composition (and mineralogy) make it difficult to constrain the extent of melting by which the pyroxenite melt was generated (i.e., for one of the pyroxenite melts, the mixing line interects it at 8% melting, while for the other it intersects it at 12% melting). For melting calculations, the following parameters were used: D for K and Ti ¼ 0.0002 and 0.008, respectively; Co for K2O and TiO2 for peridotite are 0.004 wt.% and 0.15 wt.%, respectively; for pyroxenite 1 (upper curve) 0.067 wt.% and 0.17 wt.%, respectively; and for pyroxenite 2 (lower curve) are 0.08 wt.% and 0.225 wt.%, respectively (Castillo et al., 2000) (reproduced by permission of American Geophysical Union from Geochem. Geophys. Geosys., 2000, 1, 1999GC000024).

melting that exist from region to region resulting from mantle temperature variations. Nevertheless, evidence for chemical and/or mineralogical heterogeneity of the mantle manifests itself in mantle samples recovered from ophiolites, abyssal and alpine peridotites, and mantle xenoliths (see Chapters 2.04 and 2.05), as well as in the compositions of basaltic melts (see Chapter 2.03; Natland, 1989; Salters and Dick, 2002). Indeed, major element systematics from ridge segments in the vicinity of some hotspots (e.g., the Azores, Galapagos), are anomalous with respect to the global correlations (Klein and Langmuir, 1987; Figure 10), in part due to melting in the presence of increased water in the mantle (Asimow and Langmuir, 2003). In addition, along ridge segments far from hotspots highly incompatible trace element variations are often decoupled from major element and more moderately incompatible or compatible trace element variations (Figure 12(c)); and these highly incompatible trace element enrichments commonly correlate

with enrichments in radiogenic isotope compositions (Figure 16). Lastly, the compositions of individual samples from a given region (“local” variations), as opposed to the regional averages, often show chemical systematics that differ from the global trends. Each of these provides evidence of a heterogeneous source, three of which are discussed below (pyroxenite melting, assimilation of altered oceanic crust, and local variations in ocean ridge basalt composition).

3.13.4.3.1

Pyroxenite melting

The commonly observed correlation of isotopic and trace element indices of enrichment attests to the fact that the sub-oceanic mantle is compositionally heterogeneous, even in areas distant from hotspots (see, e.g., Figure 16). One widely discussed theory suggests that the enriched-todepleted range in MORB (and near-axis seamount) compositions results from melting of enriched

Worldwide Geochemical Variations among Ocean Ridge Basalts

Figure 14 Major and minor elements versus MgO (wt.%) for representative glass inclusions in phenocrysts from axial and seamount lavas from the EPR. Each symbol represents inclusions in either plagioclase (fsp) or olivine (ol) from a particular lava. Shaded fields represent the compositional extent of fresh glass from the lavas of segments 108 300 N (light gray) and 118 200 N (medium gray); data from Batiza et al., 1996, excluding the E-MORB (K/Ti . 0.15) and FeTi basalts (FeOp . 12 wt.% and TiO2 . 2 wt.%). White boxes with symbols enclosed represent the host lava compositions. Note that the melt inclusions, with the exception of R32-5, are generally similar in composition to their host lavas (Sours-Page et al., 2002) (reproduced by permission of American Geophysical Union from Chem. Geol., 2002, 183, 237– 261).

pyroxenite veins (Hanson, 1977) and depleted peridotite matrix followed by mixing of these diverse melts to varying degrees (e.g., Prinzhofer et al., 1989; Castillo et al., 2000; Salters and Dick, 2002). It has also been suggested, for example, that these enriched veins may melt preferentially during initial melting, that they may be chemically and/or mineralogically heterogeneous, and that they may vary in size and spatial distribution (e.g., Hanson, 1977; Sleep, 1984; Zindler et al., 1984; Fornari et al., 1988; White, 1993; Lundstrom et al., 1999; Salters and Dick, 2002; Hirschmann and Stolper, 1996). Chemical differences between

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depleted (N-MORB), transitional (T-MORB) and enriched (E-MORB) ocean ridge basalts (Sun et al., 1979; Sun and McDonough, 1989) are often displayed on chondrite-normalized rare earth element diagrams (Figure 17(a)) or the so-called trace-element “spidergrams” (Figure 17(b)), in which the incompatibility of the trace elements increases from right to left (lower Ds to left). Ideally, it would be useful to distinguish the geochemical effects of source heterogeneity from those that result from variations in the extent of melting. This is a long-standing problem in studying the geochemistry of regions that exhibit differing source compositions. Although, in theory, it is possible to constrain plausible extents of melting for peridotite and vein materials, in practice, uncertainties in the composition of the enriched material can overwhelm the constraints on extents of melting. An example of this problem, using the approach of Prinzhofer et al. (1989), is shown in Figure 13 with data from the northern East Pacific Rise (EPR; after Castillo et al., 2000). Also shown are the melting trajectories of two enriched pyroxenite veins of different compositions and melts produced from the depleted peridotite matrix. It is clear from this figure that the range in primitive (low TiO2) EPR magma compositions can be modeled by mixing of depleted and enriched melts. However, because the source compositions are not well known, it is difficult to constrain with confidence the extent of melting of each lithology (although ranges in extents of melting can be plausibly estimated).

3.13.4.3.2

Assimilation of altered crust

Various types of mantle heterogeneities have been identified that differ from one another in their major element, trace element, and isotope systematics (see Chapter 2.03). A number of these distinct types of heterogeneities, particularly in isotopic composition, are most strongly manifest by hotspot volcanism (e.g., Hart and Zindler, 1989). In addition, however, it is clear that there are dispersed heterogeneities in the upper mantle that are sampled by melting at mid-ocean ridges far removed from hotspot influences. A particularly interesting class of this type of heterogeneity is one that appears to display the chemical signatures of recycled oceanic crust (Mahoney et al., 1989; Klein and Karsten, 1995; Rehka¨mper and Hofmann, 1997; Niu and Batiza, 1997; Sturm et al., 1999; Eiler et al., 2000). One location where this signature has been found is the southern Chile Ridge, an unusual setting in which an actively spreading ridge is being subducted beneath the Chile trench. Klein and Karsten (1995) showed that two of the three ridge segments nearest to the trench display, for

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Geochemistry of the Igneous Oceanic Crust

Figure 15 La and Ba (in ppm) versus K2O and TiO2 for axial and seamount melt inclusions from the EPR. The three fields represent the three different lava suites: 108 300 N (striped), 118 200 N (light gray) and seamounts (dark gray). Melt inclusions from seamount lavas are shown as filled symbols; melt inclusions from axial lavas are shown as open symbols. As a group, the melt inclusions show more diversity than the axial lava suites. The most depleted compositions are represented by seamounts melt inclusions (Sours-Page et al., 2002) (reproduced by permission of Elsevier from Chem. Geol., 2002, 183, 237 – 261).

example, low Ce/Pb values more commonly associated with the continental crust (Figure 18; Hofmann et al., 1986). These and other unusual trace element and isotopic systematics suggest the possibility that beneath this ridge segment, the ambient upper mantle has been contaminated by recycled altered oceanic crust and sediment. Similarly, Niu and Batiza (1997) argued that the recycled trace element signatures of some eastern Pacific seamounts result from mixing of subducted oceanic crust into the depleted upper mantle with subsequent upwelling, melting, and dispersal of the enriched material as mobile metasomatic fluids. As is often the case, however, the most extreme manifestation of this recycled type of heterogeneity occurs at some hotspots, where, at least in the case of the Society Islands, it may also be possible to distinguish the signatures of recycled igneous crust from those of recycled marine sediment (e.g., He´mond et al., 1994; Devey et al., 1990; Chauvel et al., 1992).

3.13.4.3.3

Local trends in basalt composition

The discussion above regarding variations in extents and pressures of melting from region to region, and the conclusion that these systematics result, to first order, from regional variations in mantle temperature is based on an examination of

the average chemical compositions of lavas erupted over significant lengths of ridge (e.g. ,100 km, depending on sampling density). It has also been recognized, however, that the compositions of individual samples from limited lengths of ridge form trends (the so-called “local trends”) that, for slow-spreading ridges, are often orthogonal to the global trends, while for fast-spreading ridges are often parallel to the global array (Klein and Langmuir, 1989; Niu and Batiza, 1993). There has been much discussion about the nature and origin of these different trends. Current evidence supports the idea that the differences in trends formed by individual samples from slow- and fastspreading ridges is due not to the difference in spreading rate per se, but rather to the composition of the mantle between densely sampled slow- and fast-spreading ridges. The difference in the local trends between slow- and fast-spreading ridges for the major element parameters Na8.0 and Fe8.0 is displayed in Figures 18 and 19. Individual samples recovered from fast-spreading Pacific ridges (Figure 20) tend to form an inverse trend within the global array. Examination of the isotopic and trace element systematics of these Pacific sample suites shows that variations in Na8.0 among the samples also correlates with indices of mantle heterogeneity, such as K2O/TiO2 (Figure 16). Thus, it appears that, while the average composition of these

Worldwide Geochemical Variations among Ocean Ridge Basalts

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larger extents of melting (or from depleted mantle higher in the melting regime) are also produced at lower pressures of melting. In this sense, the local trend for slow-spreading ridges may reflect with greater fidelity the compositions of individual (less-pooled) melt parcels produced throughout the melting regime.

3.13.4.4 Spatial Variations in Lava Compositions 3.13.4.4.1

Figure 16 K2O/TiO2 versus (a) 87Sr/86Sr, (b) 143Nd/ 144 Nd for samples from the EPR between 118 450 N and 158 N. EPR samples recovered north of 148 100 N are enclosed in a separate field (dashed field) from those recovered south of 148 100 N (solid field) (Castillo et al., 2000) (reproduced by permission of American Geophysical Union from Geochem. Geophys. Geosys., 2000, 1, 1999GC000024).

samples records the mean extents and pressures of melting, the detailed variations among individual samples from Pacific ridges reflects variations in source composition. This is in keeping with the well-accepted idea that the upper mantle beneath the EPR is heterogeneous on a variety of scales (e.g., Zindler et al., 1984; Langmuir et al., 1986; Hekinian et al., 1989; Prinzhofer et al., 1989; Reynolds et al., 1992; Niu and Batiza, 1997; Castillo et al., 2000). In contrast to fast-spreading ridges, individual samples from slow-spreading ridges tend to form trends that are orthogonal to the global array (Figure 19). This finding has been interpreted in a number of ways (e.g., Klein and Langmuir, 1989; Langmuir et al., 1992; Kinzler and Grove, 1992; Niu and Batiza, 1993). One general class of models (Klein and Langmuir, 1989) holds that the positive correlation of Na8.0 and Fe8.0 reflects the imperfect mixing of samples generated throughout the melting regime, such that melts produced by small extents of melting just above the solidus (high Na8.0) are also produced at high pressures of melting (high Fe8.0), while melts produced from

Along-axis chemical variations

With increased extent of sampling of the ocean ridge system, as well as more detailed sampling of individual ridge segments, spatial variations in axial lava compositions have been observed on a variety of scales. On the largest scale—that of an ocean basin—it has been shown that lavas erupted along much of the Indian Ocean ridge system display differences in major element, trace element, and isotopic composition compared to the majority of MORB from the Atlantic and Pacific oceans (Hamelin et al., 1986; Dupre and Allegre, 1983; see Table 2). Indian Ocean MORBs, for example, tend to have distinctive isotopic signatures with low 206Pb/204Pb and high 207 Pb/204Pb and 87Sr/86Sr (Hamelin et al., 1986; Dupre and Allegre, 1983) as well as different trace element ratios, such as Ti/Zr , 90 (Mahoney et al., 1989). There is some question about whether the distinctive Indian Ocean ridge isotopic signature may be related to the Dupal anomaly (Hart, 1984), a globe-encircling belt of anomalous ocean island isotope compositions centered at about 308 S. Large-scale differences in isotopic composition have also been noted within ocean basins, such as the slightly higher 87Sr/86Sr and 143 Nd/ 144 Nd and lower 206 Pb/ 204 Pb isotopic compositions of northern Pacific ridges compared to southern Pacific ridges (Vlastelic et al., 1999). One of the earliest recognized large-scale sources of variation in lava composition along individual lengths of ridge is the influence of nearridge hotspots. In pioneering studies of chemical variations along the northern mid-Atlantic Ridge (MAR), Schilling and co-workers showed that with increasing proximity to the Iceland and Azores plumes, over distances of 500 –1,000 km, MORBs display an increasingly pronounced trace element and isotopic signatures characteristic of each of the hotspots (Figure 21; Sun et al., 1975; Schilling et al., 1983; Schilling, 1986). The influence of other near-ridge hotspots on ocean ridge basalt compositions, and what such influences suggest about mantle and melt flow and mixing, continues to be studied at diverse locations of hotspots/ridge interactions throughout the ocean basins (e.g., Schilling et al., 1982; Hanan et al., 1986; Mahoney et al., 1989;

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Figure 17 (a) Chondrite-normalized rare earth element patterns and (b) primitive-mantle normalized trace element patterns for N-MORB, T-MORB, and E-MORB from the MAR (see caption to Table 2 for sources; normalizing values from Sun and McDonough, 1989).

Schilling et al., 1992; Kingsley and Schilling, 1998; Douglass et al., 1999; Detrick et al., 2002). Since the mid-1980s, a great deal of effort has focused on exploring smaller scales of chemical variability. Typically, these studies explore the relationship between the boundaries of discrete domains of similar chemical composition, called magmatic segments, and the various scales of tectonic segmentation of the ridge (e.g., Langmuir et al., 1986; Macdonald and Fox, 1988; Sinton et al., 1991; Batiza, 1996). Defined primarily to describe features observed along fast-spreading ridges, “first-order” tectonic segments are bounded by long-lived transform faults and large-offset overlapping spreading centers and often display significant along-axis depth

variations of hundreds to thousands of meters with central humps flanked by deeps at ridge offsets (“overlapping spreading centers” (OSCs) are ,1– 10 km offsets in the continuity of the ridge axis in which the offset limbs of the axis overlap one another). “Second-order” tectonic segments are nested within first-order segments, and are generally bounded by OSCs that typically have axial depth anomalies of a few hundred meters. Smaller scales of tectonic segmentation, “third-” or “fourth-order” segments, truncate the ridge axis into smaller, shorter-lived segments bounded by small-offsets (Macdonald et al., 1988). Increased density of sampling of the ocean system has made it possible to explore the extent

Worldwide Geochemical Variations among Ocean Ridge Basalts

Figure 18 Nb/U versus Ce/Pb for Chile Ridge glasses: segment 1 (filled circles); segment 2 (open triangles); segment 3 (open circles); segment 4 (crosses). Also shown are fields for MORB and ocean islands, marine sediments, arc lavas, altered oceanic crust, average upper continental crust (X labeled CC) and the Society and Austral Islands (which exhibit trace element and isotopic evidence of sediment and altered crust recycling) (Klein and Karsten, 1995) (reproduced by permission of Nature Publishing Group from Nature, 1995, 374, 52–57).

to which this hierarchy of tectonic segmentation also corresponds to chemically coherent “magmatic segmentation” of the ridge. Using the terminology of Sinton et al. (1991), “primary” magmatic segmentation is generally defined as lengths of ridge that erupt magmas of similar isotopic and trace element compositions, suggesting a similar mantle source composition. “Secondary” magmatic segments are defined on the basis of similarities in parental magma composition resulting from both similarities in source composition and melting processes. Finally, “tertiary” and higher-order magmatic segments may encompass lengths of ridge that display similar shallow level magma chamber and/or crustal processes (e.g., Sinton et al., 1991; Batiza, 1996). In some areas, there is clear evidence of a correspondence between the hierarchical scales of tectonic and magmatic segmentation. In an examination of lavas erupted along the southern EPR, for example, Sinton et al. (1991) showed that isotopically coherent domains (primary magmatic segments) several hundred kilometers in length are bounded by large-offset OSCs (firstorder tectonic segment boundaries). Nested within these primary magmatic segments are numerous secondary magmatic segments that share a common parental magma composition (Figure 22) and are bounded by smaller offset OSCs (second-order tectonic segments). In a detailed study of variations in lava composition along the northern EPR, Langmuir et al. (1986)

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Figure 19 Fe8.0 versus Na8.0 for individual samples from various densely sampled ridge segments. (a) Samples from “normal” ridge segments. Shaded field is for regional averages of data from “normal” ocean ridges from Figure 10. Solid triangles for Reykjanes Peninsula; Xs for 238 N to the Kane Fracture Zone; filled diamonds from the MAR 23 – 268 N; open circles from 268 N to Atlantis FZ; filled squares from Atlantis FZ to Hayes FZ; and open squares for Atlantis II FZ. Note that Atlantic data are offset to the high side of the normal field, while the Indian data are on the low side. (b) Samples from ridge segments influenced by hotspots, compared to shaded field for regional averages of hotspot-influenced ridges from Figure 10. Solid triangles for Reykjanes Peninsula as above; þ for FAMOUS and AMAR regions of the MAR. Samples from near the Oceanographer Fracture Zone (OFZ) are in three groups based on Ce/Yb as follows Ce/Yb ¼ 9.7 – 10.7 for A, ¼ 7.4– 9.1 for B, ¼ 4.2 –6.5 for C. All data have been renormalized to 100%, and corrected for interlaboratory biases. The Na8.0 and Fe8.0 algorithms for samples south of Kane and the Reykjanes Peninsula have been adjusted as in Klein and Langmuir (1989; their figure 5) (reproduced by permission of American Geophysical Union from Mantle Flow and Melt Generation at Mid-ocean Ridges, 1992, 71, 183 – 280).

showed further that small “deviations from axial linearity” (the so-called “devals,” representing third- or fourth-order tectonic segments) often corresponds with the changes in magma composition in terms of fractionation and/or enrichment (Figure 23). It should be noted, however, that even along fast-spreading ridges, where the higherorder tectonic features are most easily identified, not all large-offset transforms correspond to changes in isotopic or source composition (e.g., the Clipperton transform), and there may be variations in composition that grade continuously

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Geochemistry of the Igneous Oceanic Crust across several hundred kilometers of ridge, irrespective of the presence of OSCs or devals (e.g., 12– 148 N EPR; Castillo et al., 2000). 3.13.4.4.2

Figure 20 Fe8.0 versus Na8.0 for individual samples within each of the five transform-bounded or overlapping spreading center- (OSC-) bounded segments studied along the EPR between 118 450 N and 158 N (Castillo et al., 2000; filled circles). Also shown is the global trend for regional averages of N-MORB (from Langmuir et al., 1992). Na8.0 and Fe8.0 for EPR samples calculated as described by Castillo et al. (2000) and caption to Figure 9 (Castillo et al., 2000) (reproduced by permission of American Geophysical Union from Geochem. Geophys. Geosys., 2000, 1, 1999GC000024).

Figure 21 Latitude (8 N) versus smoothed axial depth and (La/Sm)N for the northern MAR. Latitudinal depth variations, taken at 58 intervals, are from LeDouaran and Francheteau (1981) and Vogt (1986); La/Sm data are from Schilling et al. (1983) (reproduced by permission of American Geophysical Union from Mantle Flow and Melt Generation at Mid-ocean Ridges, 1992, 71, 183– 280).

Temporal variations on lava composition

Lavas recovered from the morphologic axis of the ocean ridge system are often referred to as “zero-age” or near “zero-age” MORBs to emphasize their recent eruption. The compositions of these lavas are then interpreted as resulting from the current conditions of melt generation and evolution beneath the ridge. From this perspective, sampling of lavas off-axis, along a traverse that extends from the ridge onto older ocean crust (along a “flowline”) reveals information on the changes in lava composition erupted through time and therefore the changing conditions of melting and melt evolution. Since the early 1990s, there has been an increased interest in exploring these temporal variations in lava composition, although such studies are few compared to on-axis sampling. On a reconnaissance sampling scale, several studies have examined basalt compositions recovered by drilling at widely spaced sites on ocean crust as old as 150 Ma, primarily in the Atlantic and Indian Oceans (Keen et al., 1990; Klein, 1992; Humler et al., 1999). These studies have shown that samples recovered from old ocean crust often differ substantially in composition from zero-age basalts along the same flowline. Humler et al. (1999) showed further that the compositions of drill samples .80 Ma suggest substantially hotter sub-axial mantle producing greater extents of melting during the Mesozoic. On the flanks of the ridge axis, where lavas are still exposed above accumulating sediment, samples recovered by dredge or submersible along flowlines have revealed more detailed information on temporal variations (e.g., Reynolds et al., 1992; Perfit et al., 1994; Batiza et al., 1996). Batiza et al. (1996), for example, examined closely spaced (1 – 2 km) samples collected along three flowline traverses between 98 300 N and 118 200 N on the EPR, and extending .40 km (,800 ka) on each side of the axis. The compositions of the flowline samples were shown to be largely symmetrical about the axis and to display systematic variations in the extent of crystallization through time, suggesting regular variations in recharge and evolution of the sub-ridge AMC. With even more closely spaced sampling in the vicinity of two devals near 128 N on the EPR, Reynolds et al. (1992) showed that the along-axis length scale of chemically coherent, deval-bounded segments waxes or wanes with time. Two factors, however, complicate the straightforward interpretation that flowline sampling

Worldwide Geochemical Variations among Ocean Ridge Basalts

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Figure 22 Along-axis magmatic segmentation and sample distribution along the EPR 13– 238 S. Primary magmatic segmentation is defined on the basis of isotopic composition; secondary magmatic segmentation is defined on the basis of similar parental magma compositions. Note that most secondary magmatic segment boundaries occur at the more pronounced bathymetric depressions in the axial depth profile (Sinton et al., 1991) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 1991, 96, 6133– 6155).

Figure 23 Axial depth and various chemical parameters for zero-age basalts versus latitude for the northern EPR 5.5 – 14.58 N. Thick vertical lines delineate proposed spreading cell margins at the major OSCs and the two transform faults. Thinner vertical lines occur at small-offset OSCs and “deviations from axial linearity” (devals) along the EPR. Strontium in ppm, MgO in wt.%. Many of the deval-bounded segments form chemically distinct ridge segments (Langmuir et al., 1986) (reproduced by permission of Nature Publishing Group from Nature, 1986, 322, 422 – 429).

represents sampling of progressive temporal variations. The first factor derives from uraniumseries dating studies of basalts collected off-axis along fast- and intermediate-spreading rate ridges (e.g., Goldstein et al., 1992, 1994; Perfit et al., 1994; see Chapter 3.14). These studies revealed that samples near the axis are generally younger than would be expected based on their distance from the apparent neo-volcanic zone, suggesting that magma may erupt throughout

a relatively wide zone extending up to 4 km offaxis. (In contrast, along the slow-spreading midAtlantic ridge at 238 N, lavas from the apparently youngest volcanic edifices were dated as tens of thousands of years old (Sturm et al., 1999).) A second factor that complicates flowline studies is that detailed geologic investigations of the spatial extent of lava flows on fast- or intermediatespreading ridges show that individual eruptions may extend several kilometers both along- and

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off-axis (e.g., see Perfit and Chadwick, 1998; Sinton et al., 2002). Thus, a certain amount of overprinting of older lavas occurs during new eruptions. Indeed, investigations of the chemical variability of vertically successive lava flows recovered down-hole at ODP drill site 504B reveals a diversity of lava compositions, including variations in radiogenic isotopic composition, suggesting different source compositions (Pedersen and Furnes, 2001).

3.13.5 THE LOWER OCEANIC CRUST 3.13.5.1 Geochemical Systematics of Dikes Beneath the volcanic unit of the oceanic crust is a unit of sheeted diabase dikes believed to represent the solidified conduits of magmas derived from the melt lens that erupt to form the volcanic unit. It would therefore be expected that the dikes would have compositions similar to those of the lavas. While this is generally true, detailed comparisons of dikes and lavas in both ophiolites and the oceanic crust suggest additional complex relationships. There are relatively few studies of dikes in the in situ ocean crust, primarily because there are relatively few settings in which significant outcrops of the deeper portions of the ocean crust are exposed. One rare “tectonic window” into the lower ocean crust exists at the Hess Deep Rift, which exposes 1 Myr old crust generated at the EPR (Lonsdale, 1988; Francheteau et al., 1990, 1992; Karson et al., 1992, 2002; Gillis, 1995). Stewart et al. (2002, 2003) examined the compositions of dikes and lavas at the Hess Deep Rift. While all dikes and lavas are of basaltic composition, the majority of crystalline lavas have lower FeOT and MgO, and elevated Al2O3 and CaO concentrations compared to the dikes (Figure 24). These chemical differences and associated modal variations were attributed to the accumulation of plagioclase in the magmas that ultimately erupt as lavas on the seafloor. The accumulation of plagioclase combined with the fractionation of mafic phases lowers the magma density by more than 0.04 g cm23 relative to most dike magmas. Thus, it appears that lower-density magmas are preferentially erupted because of their increased buoyancy, resulting in the predominance of this magma type as lavas. Conversely, the majority of dikes solidified from a higherdensity magma type that is rarely represented by lavas. (This finding also underscores the importance of analyzing basaltic glass compositions in studies of basalt composition, as opposed to magma (whole rock) compositions, which represent a mixture of melt and phenocryts.) These relationships suggest that most dikes never reach the surface and erupt lava. Examination of

Figure 24 FeOT versus MgO (wt.%). crosses ¼ lavas; open diamonds ¼ dikes; filled diamonds ¼ samples from fractured basaltic outcrops not distinguishable as lavas or dikes; filled circles ¼ basaltic glass. The fractional crystallization liquid line of descent (10 Pa LLD, shown as dashed line) was calculated using MELTS v2.0.4 with starting composition of a high MgO sample, and ƒO2 buffered at QFM. Fields enclose the main composition group defined by all but three crystalline dike samples; and the low-Fe group (composed predominantly of crystalline lava samples) offset to lower MgO and FeOT and defined by samples with MgO , 7.4 wt.% and FeOT , 22 2 1.6MgO. Sample data set is available online at the American Geophysical Union electronic supplement repository. Shaded field shows composition of glasses from the EPR region at ,2.78 N reported by Lonsdale et al., 1992 (reproduced by permission of American Geophysical Union from J. Geophys. Res., 2003, 108).

phenocryst contents and modal abundances in MORBs worldwide indicates that magma density variations and crustal heterogeneity may influence magma eruptibility across a wide range of spreading rates. Specifically, lavas from slowspreading ridges have a wider range of modal phenocryst proportions compared to lavas from intermediate- or fast-spreading ridges. Presumably, lavas from slow-spreading ridges will have a greater range of densities (due to their variable phenocryst contents) and their eruption on the seafloor reflects the heterogeneous upper crust of slow-spreading ridges, where even lower crust and mantle rocks are exposed, compared to that of fast-spreading ridges. A second finding in the study of dikes exposed at Hess Deep concerns the spatial variation in their compositions (Stewart et al., 2002). Dikes were sampled over an area encompassing 25 km of an east – west flowline, representing ,3.7 £ 105 yr of crustal accretion at the EPR. Indices of fractionation (MgO), and incompatible element ratios (La/Sm, Nb/Ti) show no systematic trends along flowline. Rather, over short (,4 m) and long (,25 km) distances,

The Lower Oceanic Crust significant variations are observed in major and trace element concentrations and ratios. Modeling of these variations attests to the juxtaposition of dikes of distinct parental magma compositions. These findings, combined with studies of segmentation of the sub-axial magma chamber (AMC) and lateral magma transport in dikes along riftdominated systems suggest a model in which melts from a heterogeneous mantle feed distinct portions of a segmented axial magma reservoir. Dikes emanating from these distinct reservoirs transport magma along axis, resulting in interleaved dikes and host lavas with different evolutionary histories (Figure 25).

3.13.5.2 Geochemical and Textural Systematics of Gabbros It is generally agreed that the thick gabbroic unit at the base of the oceanic crust forms by the relatively slow crystallization of basaltic magma at depth. Decades ago, it was envisioned that crystallization along the walls and floor of a large, predominately molten, steady-state magma chamber at depth beneath the ridge would produce the gabbroic layer and, in particular, account for

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the various types of mineralogic and chemical layering observed in many ophiolite complexes (e.g., Pallister and Hopson, 1981). With improved seismic and tomographic imaging of the region beneath the crust (e.g., Detrick et al., 1987; Toomey et al., 1990; Caress et al., 1992), it has become clear that no large, predominately molten magma body exists beneath the ridge, but rather, at least at fast-spreading ridges, a thin (,50 m), narrow (0.5 –1.5 km) melt lens at 1 –2 km depth beneath the ocean floor caps a wider region of low seismic velocities, suggesting the presence of melt and significant amounts of crystals (Figure 3). This region, commonly called the “mush zone,” may extend to 10 km in width and the fraction of melt within it is believed to decrease both with depth and laterally, ultimately grading laterally into seismic velocities consistent with solidified gabbro (e.g., Sinton and Detrick, 1992). As discussed by Coogan et al. (2002), there are currently two classes of models for the formation of the gabbroic sequence that are, to first order, consistent with available constraints from marine seismic investigations, and structural and petrologic studies of gabbros in ophiolites and oceanic drill cores. In the first class of models (Figure 25(a)), crystallization occurs

Figure 25 Model for construction of the upper crust at fast-spreading ridges (a), leading to the cross-sectional (flowline) variations in dike compositions (b) exposed on the north wall of the Hess Deep Rift. The melt lens is fed by distinct mantle melts along axis. Incomplete mixing along the length of the AMC for the equatorial EPR produces compositionally distinct portions of the AMC (distinct shading). Dikes are injected in narrow regions near the axis (1 – 2 km wide), but incomplete mixing in the AMC and lateral (along-axis) magma transport (shown as vectors in the dikes) results in closely spaced dikes of distinct compositions and the absence of a simple trend in composition along a flowline section. Dikes will tend to pierce the surface and erupt lavas over their initial site of injection, but many dikes will not intersect the surface and erupt lavas (see text for discussion). Note that in panels (a) and (b), the dike thickness (commonly ,1 m) has been exaggerated for visual clarity and the top of the melt lens has been placed at 2 km depth, but ranges from ,1 – 2 km (Stewart et al., 2002) (reproduced by permission of American Geophysical Union from J. Geophys. Res., 2002, 107, 2238).

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Figure 26 Schematic drawing for recently proposed crustal models. (a) Ductile flow model in its simplest form (e.g., Phipps Morgan and Chen, 1993; Quick and Denlinger, 1993; Nicolas et al., 1988). Ductile flow down and outward from a mid-crustal magma chamber constructs the lower crust. (b) Ductile flow model incorporating a basalt sill (Schouten and Denham, 1995). (c) Hybrid ductile flow model with on- and off-axis sill intrusions (Boudier et al., 1996). (d) Sheeted sill model with melt transport by hydrofracture and in situ emplacement of the lower crust by on-axis sill intrusions (Kelemen et al., 1997; Bedard et al., 1988) (reproduced by permission of Elsevier from Planet. Sci. Lett., 1998, 156, 1 – 11).

predominantly within the melt lens (and possibly a second melt lens at the base of the crust; Figure 26(b)). The resulting gabbroic material subsides and/or is transported laterally off-axis (e.g., Sleep, 1974; Dewey and Kidd, 1977; Nicolas et al., 1988; Quick and Denlinger, 1993; Phipps Morgan and Chen, 1993). Evidence for this comes primarily from textural studies of gabbroic rocks in the Oman ophiolite that show evidence of solid-state ductile deformation in the absence of significant crystal – plastic deformation, suggesting that large strains may have been accommodated by deformation of a crystal mush. In the second class of models (Figure 26(d)), gabbro formation occurs primarily within the mush zone in dispersed sills (Korenaga and Kelemen, 1997, 1998; Kelemen et al., 1997). Evidence supporting this latter model for the Oman ophiolite includes the observation that discrete gabbro sills in the crust –mantle transition zone have textures and chemical compositions similar to those of the modally layered gabbros of the thick, lower gabbro section. In addition, minerals of the lower gabbros are in chemical equilibrium with melt compositions of the upper gabbros, dikes and lavas, suggesting that the latter three formed in large part from melts expelled from the crystallizing lower gabbro sills (Kelemen et al., 1997). Structural and other studies of Oman gabbros support the possibility that both of these processes may be occurring simultaneously (Kelemen and Koregan, 1997), a conclusion also reached by Coogan et al. (2002) in a geochemical and textural study of gabbros exposed at Hess

Deep. These two models, however, were primarily developed for fast-spreading ridges, and as emphasized by Natland and Dick (2001) may not be applicable to gabbros formed at slowspreading ridges, such as the petrologically complex gabbros recovered in ODP Hole 735B in the Indian Ocean. 3.13.6 CONCLUSIONS Much is known about the origin of the chemical variability of the oceanic crust, but much also remains enigmatic. At present, we have relatively sophisticated theories on key aspects of the processes and variables involved in the generation, evolution and emplacement of magma to form the crust. But the linkages between these processes remain speculative. One example of this concerns the enigmatic relationship between the diverse melt compositions produced in the mantle and the various scales of tectonic and magmatic segmentation observed along the ridge axis. Are some of the scales of tectonic and magmatic segmentation a reflection of deep mantle upwelling processes while others merely reflect shallow magmatic plumbing systems? Furthermore, while we have an extensive catalog of the diverse types of heterogeneities in the mantle, no unifying paradigm has emerged to fully explain their relationships to one another, how they developed, and their spatial scales. Lastly, combined geochemical, geophysical, and geological studies are needed to elucidate the processes of crustal accretion that produce the architecture of the ocean crust.

References ACKNOWLEDGMENTS I am grateful to Roberta Rudnick for her insightful editing and unflagging encouragement, and to Terry Plank for her careful review of this manuscript. This work was supported by NSF Grant #OCE- 0099154.

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Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 433–463

3.14 Melt Migration in Oceanic Crustal Production: A U-series Perspective T. Elliott University of Bristol, UK and M. Spiegelman Columbia University, New York, NY, USA 3.14.1 INTRODUCTION 3.14.2 U-SERIES PRELIMINARIES 3.14.2.1 Naturally Occurring Actinide Decay Chains 3.14.2.2 Closed-system Models 3.14.2.2.1 Bulk system models 3.14.2.2.2 Multiphase models 3.14.2.3 Measurement and Nomenclature 3.14.3 OBSERVATIONS 3.14.3.1 Data Compilation 3.14.3.2 Age Constraints 3.14.3.3 General Signatures of Disequilibrium 3.14.3.4 Assessment of Shallow-level Contamination 3.14.3.5 Further Observations 3.14.3.5.1 Equiline diagram 3.14.3.5.2 Source enrichment 3.14.3.5.3 Variations of disequilibria with spreading rate 3.14.3.5.4 Variations of disequilibria with axial depth 3.14.3.5.5 Mutual covariations of disequilibria 3.14.3.6 Summary of Observations 3.14.4 U-SERIES MELTING MODELS 3.14.4.1 Preliminaries: Partition Coefficients and Diffusivity 3.14.4.1.1 Diffusion coefficients for U-series nuclides 3.14.4.2 Implications for Closed-system Models 3.14.4.3 Open-system “Ingrowth Models” 3.14.4.3.1 “Dynamic melting models” 3.14.4.3.2 Single porosity transport models 3.14.4.3.3 Comparison with observations 3.14.4.3.4 Two-porosity models 3.14.4.3.5 Full reactive transport models 3.14.5 SUMMARY OF MODEL BEHAVIOR 3.14.6 CONCLUDING REMARKS ACKNOWLEDGMENTS REFERENCES

465

466 466 466 467 467 468 469 472 472 472 475 477 479 479 479 484 484 487 489 489 489 491 491 492 492 494 497 500 501 504 505 506 507

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3.14.1 INTRODUCTION The crust is ultimately the product of mantle melting. To understand its formation requires knowledge of how melts move from depth. Generation of oceanic crust as a consequence of seafloor spreading is clearly related to melt production below. Although melting processes beneath ridges are perhaps the best constrained of any tectonic setting (see Chapter 3.13; Langmuir et al., 1992), neither melt distributions nor rates of melt migration are well resolved. It is necessary to infer dynamic processes hidden within the mantle by using proxy observations. Seismology provides important but static information on the velocity structure of the mantle, which is sensitive to the gross amount and distribution of partial melts (e.g., Forsyth et al., 1998). The geochemistry of erupted melts and their residues provides a different perspective, since their compositions are sensitive to the pathways they follow (Hellebrand et al., 2002; Johnson et al., 1990; Kelemen et al., 1995a; Klein and Langmuir, 1987; Salters and Hart, 1989; Spiegelman, 1996). Furthermore, short-lived radioactive nuclides can provide crucial information on the rates of these processes. The aim of this chapter is to review the observations of uranium series nuclide studies on mid-ocean ridge basalts (MORBs) and discuss their implications for melt transport processes. We note that a recent review by Lundstrom (2003) has a similar remit and makes a useful companion contribution. We initially recap the fundamental behavior of the naturally occurring actinide decay chains, and outline the analytical challenges of

238

Uranium

234

U

4.47×10 9 yr

3.14.2 U-SERIES PRELIMINARIES Ivanovich and Harmon (1992) and Bourdon et al. (2003) provide comprehensive reviews of the fundamental behavior of U-series systems. Here we only briefly reiterate these properties, focusing on those relevant to melt migration from the mantle.

3.14.2.1 Naturally Occurring Actinide Decay Chains The major uranium nuclides 238U and 235U decay to the stable nuclides 206Pb and 207Pb, respectively, through a long chain of intermediate daughter products (Figure 1). The half-lives of 238U and 235U are ,4.5 Gyr and 0.7 Gyr, respectively. All intermediate nuclides have halflives much shorter than the uranium nuclides at the head of the chain, but still span a huge range of timescales from ,245 kyr to 164 ms. The longerlived intermediates have half-lives appropriate

235

U

U

7.04×10 8 yr

2,45,000 yr 234

Protactinium

their measurement. We then evaluate the gross signatures of disequilibrium observed in MORBs and consider the co-variations of disequilibria with other geochemical and geophysical parameters. This is followed by a detailed discussion of models of melt transport that have been developed to account for these observations, allowing a synthesis of the key constraints that uranium-series (henceforth U-series) measurements provide on the process of melt migration beneath mid-oceanic ridges.

231

Pa

Pa

32,800yr 234

Thorium

Th

230

231

Th

a-decay

75,500 yr

Th 227

Actinium 226

Radium

b-decay

227

Th

223

Ra

219

Rn

Ac

22.1 yr

Ra

1,600 yr 223

Francium 222

Radon

Rn 218

Astatine 218

Polonium

214

Po

Bismuth Lead

219

At

214 214

Fr

Po 210

Bi 210

Pb

210

215 206

Pb

215

Po

Bi Pb

215

At

At

Po 211

Bi 211

211

Po

207

Pb

Bi

Pb

22.3 yr 210

Thallium

Figure 1

238

U and

235

Tl

206

Tl

207

Tl

U decay chains. Longer-lived nuclei discussed in the text are shown with their half-lives (in years) listed beneath.

U-series Preliminaries for studying melt migration. The principal nuclides of interest are 230Th ðt1=2 , 75 kyrÞ and 226 Ra ðt1=2 , 1:6 kyrÞ in the 238U chain and 231Pa ðt1=2 , 33 kyrÞ in the 235U chain (Figure 1). The longest lived daughter nuclide is actually 234U ðt1=2 , 245 kyrÞ; but this is neither expected (Fleischer et al., 1975) nor observed (see Section 3.14.3.3) to fractionate from 238U in high-temperature igneous systems. Thus, 234U measurements do not provide temporal information, but do provide a useful screen for low temperature alteration (see Section 3.14.3.4). A third naturally occurring actinide decay chain is headed by 232Th and decays to 208Pb. The intermediate nuclides in this chain have not been used to elucidate melting processes beneath ocean ridges, and here we only concern ourselves with the behavior of the 232Th parent. 232Th is the overwhelmingly dominant Th nuclide, and is important as a reference nuclide for the shortlived 230Th nuclide in the 238U chain. Although the decay of 238U, 235U, and 232Th support chains of short-lived daughter nuclides, the effect of this decay in changing the abundances of the parent nuclides themselves is nevertheless insignificant on the 100 kyr timescales considered here. Thus, 238U, 235U, and 232Th can be paradoxically considered to behave as “stable” reference nuclides on the timescales considered here. 3.14.2.2 Closed-system Models To understand the behavior of U-series nuclides, we start by discussing closed systems. These models form the basis for more complicated open-system models discussed later. Here we discuss the behavior of generic radioactive decay chains that apply to all the naturally occurring actinide chains. 3.14.2.2.1

Bulk system models

The equations and solutions for closed-system radioactive decay chains have been known since Bateman (1910). To understand the behavior of these systems, however, it is useful to express them as a linear system of ordinary differential equations and use some basic results from linear algebra to discuss the general solutions. This treatment helps to elucidate the ideas of “secular equilibrium” and relaxation to equilibrium. Consider a general decay chain of nuclides such that element 1 decays to element 2 decays to element 3, etc. (e.g., 238 U ! 230 Th ! 226 Ra). In discussing U-series nuclides, it is convenient to represent their abundances in terms of activities. The activity of a nuclide is its rate of decay, which

467

depends only on the number of atoms of the nuclide and the decay constant l: Thus, for a chain of nuclides, we can write the (specific) activity of the ith element in the chain ði ¼ 1; 2; 3; …Þ as ai ¼ li ci ; where c is atomic concentration by weight. The evolution of activities in a decay chain of length n is then described by da1 dt da2 dt da2 dt .. . dan dt

¼ 2l1 a1 ¼ l2 a1 2 l2 a2 ¼ l3 a2 2 l3 a3

ð1Þ

¼ ln an21 2 ln an

For every nuclide (after the parent), the change in activity depends on the difference between decay of the daughter and production by the parent. Using linear algebra, Equation (1) can be written as a vector equation of activities 32 3 2 3 2 2l1 a1 a1 76 7 6 7 6 76 7 6 7 6 76 a2 7 6 a2 7 6 l2 2l2 76 7 6 7 6 7 6 76 7 d6 76 a3 7 6 a3 7 6 l 2 l 3 3 76 7 6 7¼6 7 6 76 7 dt 6 76 . 7 6. 7 6 . . 76 .. 7 6 .. 7 6 . . . . 76 7 6 7 6 54 5 4 5 4 ln 2ln an an ð2Þ or more compactly using matrix-vector notation as da ¼An a dt

ð3Þ

where An is the matrix of decay constants in Equation (2). If all the decay constants are distinct then Equation (3) can be solved analytically using standard approaches of linear algebra (e.g., Strang, 1998) as aðtÞ¼SeLt S21 a0

ð4Þ

where S is the matrix of eigenvectors of An, L is a diagonal matrix of eigenvalues of An ; which happen to be the diagonal elements of An (i.e., 2l1 ;2l2 ;…;2ln ) and a0 is the initial activities at time t ¼ 0: For this problem, both S and L can be found analytically, although for more general problems they can also be found numerically with packages such as MATLAB. (A caution though. For numerical stability, it is important to solve the equations for the activities rather than the concentrations because the matrix that arises from the concentration equation is ill-conditioned if the decay constants are wildly

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different.) Equation (4) can be written more physically as aðtÞ¼ c1 s1 e2l1 t þc2 s2 e2l2 t þ··· þcn sn e2ln t

ð5Þ

where si is the ith eigenvector and ðc1 ; c2 ;…; cn Þ¼ c¼S21 a0 is the decomposition of the initial condition into its component eigenvectors. Physically, Equation (5) states that the system can be considered composed of n independent vectors that each decay like one of the nuclides. These eigenvectors are the supported chains, for example, in the series 238 U! 230 Th! 226 Ra; any initial set of activities can be decomposed into three chains, one supported by 238U that decays slowly, one supported by 230Th that decays with a half-life of ,75 kyr and one supported by 226Ra that decays rapidly. After a sufficiently long time, the shorter-lived eigenvectors decay away leaving just the most slowly decaying one. This is the state of secular equilibrium in which all intermediate nuclides are supported by the parent nuclide at the head of the chain because this is the slowestdecaying species. In addition, if the decay constants are very different in magnitude (e.g., l1 p l2 p l3 ) then the activities in any given supported chain will be nearly identical such that their activity ratios ðai =aj Þ are ,1. The statement that all the ratios of all activities are 1 in secular equilibrium is a convenient short-hand notation for the longer-lived U-series nuclides. Any disturbance from secular equilibrium decays back toward secular equilibrium through the decay of the shorter-lived supported chains. If the chains have significantly different decay rates, they are reasonably well decoupled and can be used to date processes comparable to the lifetimes of each chain (e.g., Condomines et al., 1988; Rubin et al., 1994; Thompson et al., 2003). As a useful rule of thumb, the time taken for a parent – daughter pair to return to approximate secular equilibrium, is about five half-lives of the shorter lived nuclide. In five half-lives, ,97% of initial disequilibrium has decayed. Whether any detectable disequilibrium actually remains depends on the precision of measurements and degree of initial disequilibrium. 3.14.2.2.2

Multiphase models

The closed-system model is also used to justify the assumption that the solid mantle source is in secular equilibrium before melting begins. This will be the case if the mantle has acted as a closed system for more than about five half-lives of the longest-lived daughter product perturbed in the most recent event. As mentioned above, 234U is not fractionated from 238U at mantle temperatures and so 230Th ðt1=2 , 75 kyrÞ constrains the time

for return of the mantle to secular equilibrium (in a closed system) to be only some 400 kyr. This bulk state of secular equilibrium applies to the total amount of the U-series nuclides, but does not necessarily say where the different elements reside within the system. If the bulk system has a single phase (such as a melt or a monomineralic rock) then that phase will be in secular equilibrium. If the material has multiple phases with different partitioning properties, however, the individual phases can maintain radioactive disequilibria even when the total system is in secular equilibrium. There are two basic sets of models that exploit this fact, the first assumes complete chemical equilibrium between all phases and the second assumes transient diffusion controlled solid exchange. (i) Equilibrium models. In the equilibrium multiphase models, the closed system is assumed to be in both chemical and secular equilibrium at all times (e.g., Alle`gre and Condomines, 1982; Condomines et al., 1988; Condomines and Sigmarsson, 2000). For example, if we have a rock with N solid phases plus melt where each phases occupies Xi mass fraction and has a melt/mineral equilibrium partition coefficient Di ; then the total activity ratio of (230Th/238U) ¼ 1 implies that ! 230 Th ¼1 238 U total X  230 N Th X D þ X Th f f i¼1 i i ð6Þ  ¼ X 238 N U X D þ Xf Uf i¼1 i i where (230Th)f and (238U)f are the activities of Th and 238U in the melt and Xf is the mass fraction of melt. Rearranging shows that the activity ratio in the melt is ! 230 Th ð1 2 Xf ÞDU þ Xf ¼ ð7Þ 238 U ð1 2 Xf ÞDTh þ Xf f

230

where DU and DTh are the bulk melt/solid partition coefficients for U and Th, respectively. Thus, the melt can support activity ratios different from 1 if the melt fraction Xf is smaller than the partition coefficients. In the limit of infinitesimal melt fractions, the largest excess is approximately DU/DTh. While there is still some debate about the bulk partition coefficients of U-series nuclides in mantle systems (see Section 3.14.4.1), they are all uniformly small ð&0:005Þ: Thus, pure chemical fractionation due to equilibrium batch melting can produce melts with disequilibrium activity ratios but only for very small degrees of melting (i.e., Xf , DU) (e.g., Condomines et al., 1988; Condomines and Sigmarsson, 2000).

U-series Preliminaries These small degree melts should also be extremely enriched in highly incompatible elements and would resemble alkali basalts rather than MORBs (unless the source was extremely depleted). This mechanism cannot be ruled out for some ocean island basalts (OIBs) (Sims et al., 1999; Elliott, 1997). However, it will not work for the significant degrees of melting ðF * 10%Þ inferred for MORBs (Klein and Langmuir, 1987; McKenzie and Bickle, 1988; see Chapter 3.13). Moreover, to preserve the excesses produced by this mechanism requires preferential extraction of the enriched phase (e.g., the melt) and extraction on a timescale that is short compared to shortest-lived nuclide of interest. Nevertheless, this basic mechanism of using specific phases to concentrate nuclides chemically, followed by preferential extraction of those phases, is a process that is common to many of the models proposed for U-series excesses including the “dynamic melting models” (McKenzie, 1985; Williams and Gill, 1989) discussed in Section 3.14.4.3.1 as well as more recently proposed solid state, diffusion controlled models (e.g., Van Orman et al., 2002a; Saal et al., 2002b; Feineman et al., 2002) (ii) Diffusion limited models. Chemical partitioning into specific phases is easily extended to multiphase solids that contain no liquid. For example, Van Orman et al. (2002a) and Saal et al. (2002a,b) suggest that partitioning of radium into plagioclase in the crust with subsequent re-melting of the plagioclase could be a possible source of radium excesses in MORBs. A similar model is proposed for arcs where the principal phase is assumed to be phlogopite (Feineman et al., 2002). These models allow radioactive disequilibrium in individual solid phases to be preserved in steady state. This arises due to the competing effects of chemical and secular equilibrium between phases in certain mineral assemblages. Three key conditions are required: First the ratios of melt/solid partition coefficients for a parent and daughter nuclide pair are greatly different in two coexisting phases. Second, the solid diffusivity of at least one of the parent or daughter nuclides must be sufficiently large so that the most diffusive element is mobile on a timescale comparable to the decay time of the shortest-lived nuclide. Third, there must be sufficient time for diffusion and ingrowth to take place. These models are reasonably new and need to be tested against their implications for other elements. 3.14.2.3 Measurement and Nomenclature Activities of U-series nuclides in MORBs have been determined by both nuclear particle counting and mass-spectrometric techniques. In this section, we highlight briefly the key

469

differences in the analytical methods and discuss the associated nomenclature to avoid any confusion such diversity might cause. U-series nuclide activities can be measured directly by detection of their emitted nuclear particles, e.g., alpha particle counting by solidstate detectors (Ivanovich and Harmon, 1992). In contrast, measurements by mass-spectrometry do not require waiting for Nature to take its course. Atoms of the sample are ionized and accelerated so that charged particles of the nuclides themselves can be measured by Faraday cups or electron multipliers (see Goldstein and Stirling, 2003). Mass-spectrometry is hence a more rapid technique. Typically mass-spectrometry measurements take tens of minutes to hours, while counting methods require days to weeks. The precision of the techniques ultimately depend on the number of nuclear decays (particle counting) or ionized particles detected (massspectrometry). Both are related to the number of atoms of the nuclides of interest present in the sample, but with a different functionality. As discussed above, the natural decay rate is the product of the number of atoms present and the activity constant (the reciprocal of half-life multiplied by the natural log of 2, l ¼ ln 2=t1=2 ). The activities of all nuclides in the U-series chain are equal at secular equilibrium, and generally within a factor of 3, except in cases of extreme disequilibrium (see Section 3.14.3.3). Thus all U-series nuclides can potentially be measured with similar precision using nuclear counting techniques. The number of charged particles produced by mass-spectrometry is the product of the number of atoms of a nuclide multiplied by its ionization efficiency. The number of atoms of a nuclide equals the nuclide’s activity divided by its activity constant. Since the activities of all U-series nuclides are generally similar but activity constants vary by orders of magnitude, the atomic abundances of the U-series nuclides also vary by orders of magnitude. Longer-lived nuclides (with smaller activity coefficients) are present in higher atomic abundances than the shorter-lived nuclides, and so are likely more amenable to analysis by mass-spectrometry. The ionization efficiencies of the U-series nuclides vary by orders of magnitude and so this too must be considered (Figure 2). “Traditional” thermal ionization mass-spectrometry efficiently ionizes Ra (.1%) (Cohen and Onions, 1991; Volpe et al., 1991). Elements with higher first ionization energies, such as Th and U, give significantly lower ion yields, typically ,1‰ (see Edwards et al., 1987) although work by Yokoyama et al. (2001) reported ionization efficiencies up 5‰ for U. Yet, even given typical thermal ionization efficiencies of 0.5– 0.03‰ for 10 –100 ng Th loads

470

Melt Migration in Oceanic Crustal Production 0.25 alpha (2g sample, 2 weeks counting) mass-spec (0.1 g sample, ion. eff. 0.00005) mass-spec (0.1 g sample, ion. eff. 0.0005) mass-spec (0.1 g sample, ion. eff. 0.01)

2s counting error

0.20

0.15

0.10

0.05

0.00 1,000

10,000

226

Ra

100,000

230

Th Measured nuclide half-life (yr)

1000,000

234

U

Figure 2 Variations of counting statistical error (2s or 95% confidence) for nuclides of different half-lives using different measurement techniques. Four different scenarios are shown with details listed in legend, aimed to cover a typical range of conditions. The values of ionization efficiencies span a full range of values appropriate for elements difficult (Th) and easy (Ra) to ionize thermally, see text. All scenarios assume a sample with MORB-like U concentration (50 ng g21), but with all daughter nuclides in secular equilibrium (for illustrative simplicity). Calculations assume (unrealistic) 100% yields for chemical purification of the nuclide of interest and 40% counting efficiency for alpha detection geometry.

(Asmerom and Edwards, 1995; Edwards et al., 1987; Goldstein et al., 1989; Lundstrom et al., 1999; Rubin, 2001), the total ion yield for the longer-lived nuclides of interest here is significantly greater than the number of nuclear decays over reasonable counting times. The corollary of this is, of course, better analytical precision (Figure 2). There are no alpha counting analyses of MORB for the decay chain of 235U, which is ,138 £ less abundant than 238U. Yet, mass-spectrometry has provided a valuable data set of (231Pa/235U) disequilibrium in MORBs (Goldstein et al., 1993; Lundstrom et al., 1995; Sims et al., 2002). Furthermore, ion yields of 2– 5‰ for Th and U have recently been achieved by plasma-ionization mass-spectrometry (Luo et al., 1997; Pietruszka et al., 2002; Turner et al., 2001) and efficiencies as high as ,2% reported for secondary ion mass-spectrometry (England et al., 1992; Layne and Sims, 2000). Such advances should boost the precision of mass-spectrometry (e.g., Pietruszka et al., 2002). The counting statistical limit on precision is of importance to studies of MORBs due to the generally low abundances of uranium and its daughter nuclides. Naturally, counting statistics are only a

best case limit on the precision and accuracy of analyses. A number of additional chemical preparation (e.g., blanks, sample-spike equilibration) and instrumental problems (e.g., detector intercalibration, interferences) can considerably degrade data quality. As discussed by Rubin (2001), comprehensive assessment of data quality is difficult without associated publication of detailed analytical procedures, reproducibility of solution standards and full replicates of samples and international rocks standards with similar matrices and nuclide abundances. Not all MORB studies, whether by particle counting or massspectrometry, provide such complete information. But counting statistics or internal precisions are commonly reported. Whilst these data are not robust assessments of total error, it is important to note that data cannot be more accurate than they are precise, and so they provide minimum constraint on uncertainties. Given the large variability of actinide abundances in MORBs (e.g., Figure 3) and different laboratory procedures, it is difficult to make definitive statements about the precision obtainable by different techniques. As useful rule of thumb, however, 238U – 230Th – 226Ra and 235U – 231Pa

U-series Preliminaries

471

3.0 alpha (Newman et al., 1983) alpha (Reinitz and Turekian, 1989) alpha (Rubin and Macdougall, 1988) mass-spec

( 238 U/ 232 Th)

2.5

2.0

1.5

1.0

0.5 0.01

0.10

1.00

10.00

[Th] (µg g–1)

Figure 3 Concentration of Th (mg g21) versus (238U/232Th) for compiled MORB data set. Note logarithmic scale for [Th]. Mass-spectrometric measurements in filled black symbols, cited errors not shown but typically ð2sÞ precision comparable or smaller than sample size. Studies with (238U/232Th) measured by alpha counting shown in open symbols and identified individually in legend. Cited counting errors ð2sÞ shown for (238U/232Th) alpha measurements, but not for [Th] which are comparable to sample size. Note some of the alpha counted samples extend to values of (238U/232Th) much higher than the array defined by mass-spectrometry data.

disequilibria in MORBs are measured with an internal precision (95% confidence or 2s) of typically ,0.5 –2% by mass-spectrometry compared to counting statistical errors of 5– 15% (2s) for the 238U chain disequilibria by alpha counting (Figure 2). U-series disequilibria are most naturally expressed in terms of activity ratios (e.g., Section 3.14.2.2). Alpha counting measures activity directly, whereas mass-spectrometry yields atomic ratios which need to be converted into activities using activity constants. This introduces an additional component of uncertainty (1 – 8‰) to the absolute accuracy of mass-spectrometric activity measurements (e.g., Holden, 1989; Jaffey et al., 1971; Meadows et al., 1980). This uncertainty, however, is small compared to the uncertainties in particle counting measurements. Moreover, the high precision of massspectrometric measurements has allowed some activity constants to be refined using samples where secular equilibrium can be assumed (Cheng et al., 2000). Owing to the different measurement techniques, U-series nuclide concentration measurements are reported in a number of ways. Specific activities of each nuclide (e.g., dpm g21) are reported

for alpha counted studies, whereas mass-spectrometric measurements provide weight concentrations of nuclides (e.g., mg g21) and atomic isotope ratios. As discussed above, disequilibria are expressed in terms of isotope activity ratios, conventionally enclosed in curved brackets to distinguish them from weight ratios (commonly reported in square brackets) and atomic ratios (with no brackets). As well as activity ratios, disequilibria are sometimes expressed as “excesses” or “deficits.” An excess is the fraction of unsupported daughter nuclide, e.g., (230Th/238U) ¼ 1.2 can be described as a 20% 230 Th-excess. The (234U/238U) ratio is also sometimes reported in a d-notation or parts per thousand relative deviation from equilibrium. Concentrations of uranium and thorium are also measured by other techniques (e.g., instrumental neutron activation analysis) and sometimes included in trace element analyses. In geochemical literature the ratio [Th/U] is frequently used. However, 232Th is the only long-lived thorium isotope and 238U is a fixed 99.28% of total uranium, so [Th/U] ratios are closely related (3.3% smaller) to atomic 232Th/238U ratios. The latter ratio has historically been termed k when used in the Pb isotope literature. To further

472

Melt Migration in Oceanic Crustal Production

the confusion, the activity ratio (238U/232Th), is used in the common equiline plot. Although all simply interrelated, the various different expressions of Th/U ratios can seem puzzling.

3.14.3 OBSERVATIONS 3.14.3.1 Data Compilation Since we wish to examine the implications of U-series disequilibria on processes of melting and melt migration, it is necessary to compile a dataset of samples with (near) “primary” mantle signatures. Inevitably, decay of disequilibria after eruption will modify primary values. Indeed, a number of U-series studies of MORBs have specifically targeted off-axis samples and used disequilibrium as a high-resolution chronometer of oceanic plate accretion (e.g., Goldstein et al., 1991; Krishnaswami et al., 1984; Sims et al., in press; Sturm et al., 2000; Zou et al., 2002). For our purposes, however, it is essential to filter out samples from the literature which have likely experienced significant post-eruptive decay. In the following section we describe age constraints on young MORBs, which we have used to guide data compilation for this contribution (Table 1). A second issue in sample selection is possible crustal or seawater contamination (Section 3.14.3.2). We include in our compilation only analyses from picked, fresh glass. Hand picking of glasses critically allows visual rejection of contaminated material (see Bourdon et al., 2000). We have excluded samples where authors have themselves inferred contamination (e.g., Condomines et al., 1981; Goldstein et al., 1991; Newman et al., 1983), but no further screening has been applied. Additional sample selection criteria have been applied to try to exclude some samples with anomalous ( 238U/ 232 Th) and ( 230Th/ 232Th) ratios. Figure 3 illustrates a plot of [Th] against (238U/232Th) ratios for MORBs from this compilation. Some of the alpha-counted measurements are clearly elevated relative to the general array defined largely by more precise massspectrometric determinations of uranium and thorium concentrations. Reinitz and Turekian (1989) compared the high (238U/232Th) obtained by themselves and Rubin and Macdougall (1988) to earlier analyses from a similar location by Newman et al. (1983) and concluded that analytical differences most likely explained the differences between the contrasting data sets. Additional samples measured from ,128 N on the EPR (Reinitz and Turekian, 1989; Rubin and Macdougall, 1988) also have much higher

(238U/232Th) than later isotope dilution massspectrometric measurements from the same area (Ben Othman and Alle`gre, 1990). New massspectrometric measurements of Rubin et al. (2000) have demonstrated some of his alpha counting measurements on low abundance uranium samples gave inaccurate (238U/232Th) and (230Th/232Th) due to blank problems (Rubin, personal communication). It is clearly difficult to assess definitively which samples may have been affected. To avoid any of the anomalous samples evident in Figure 3, we exclude from later discussion (Section 3.14.3.5) samples with [Th] , 0.5 ppm measured by Rubin and Macdougall (1988) and Reinitz and Turekian (1989). This is an arbitrary cut-off, which may exclude accurate data, but given the known problem, it seems an appropriate approach until further analyses are published. Nevertheless, we include all these samples in the initial general discussion of disequilibria (Section 3.14.3.3), as their (230Th/238U) are not obviously anomalous. We consider exclusively MORB samples, and not OIBs that compromise a smaller portion of the oceanic crust. Both MORBs and OIBs are believed to be derived from decompression melting of the mantle. A comparison of the magma types is interesting but beyond the scope of this contribution; see Bourdon and Sims (2003) for a recent review of the U-series systematics of OIBs. Nevertheless, we do include samples from elevated ridge segments believed to be influenced by OIB-source mantle. In these cases, however, the melting regime should fundamentally resemble that of MORBs. Potentially, a component of active upwelling further complicates interpretation of U-series results from OIBs influenced segments (Bourdon et al., 1996b), as has been discussed for OIBs (e.g., Chabaux and Alle`gre, 1994; Sims et al., 1999). Yet work on sub-aerial magmatism at the most northerly end of the Icelandic rift suggests that there is little active upwelling this far from the putative plume center (Maclennan et al., 2001). This conclusion also likely applies for the Reykjanes Ridge and Azores platform samples considered here, which lie even further from their respective centers of sub-aerial volcanism. 3.14.3.2 Age Constraints The effects of post-eruptive decay are illustrated in Figure 4, using disequilibria typical for MORBs (see Section 3.14.3.3) as initial conditions. The problem of recovering samples with “primary” signatures is most acute for (226Ra/230Th). Any 226Ra-excesses will be only 1% of their initial values 10 kyr after eruption. Over the same time, the change in typical (230Th/238U) is barely resolvable analytically, but

Table 1 Location

Lat

Summary of axial MORB samples used in this compilation, grouped by geographical location.

Depth (m)

Half-spreading rate (cm yr21)

Juan de Fuca

44– 488 N

2,200

3

Gorda

41– 428 N

3,000

EPR

20– 218 N

EPR

(230Th / 238U)

(231Pa / 235U)

(226Ra / 230Th)

Equiline gradient

err

Reference

Goldstein et al., 1991, 1993, Volpe and Goldstein, 1993, Lundstrom et al., 1995 Goldstein et al., 1991, 1993; Volpe and Goldstein, 1993; Cooper et al., 2003 Newman et al., 1983; Rubin and Macdougall, 1988, 1990 Rubin and Macdougall, 1988; Reinitz and Turekian, 1989, BenOthman and Alle`gre 1990; Goldstein et al., 1991, 1993 Volpe and Goldstein, 1993; Lundstrom et al., 1999; Sims et al., 2002 Rubin and Macdougall, 1988, 1990 Rubin and Macdougall, 1988, 1990

13

5

5

0.37 (0.48)

0.05

2.75

9

4

4

0.27

0.05

2,600

3

6

1

20.17

0.33

11– 138 N

2,600

5.5

6

10

0.053

0.18

EPR (þSiqueiros transform)

8 – 108 N

2,600

5.5

40

23

0.49 (0.66)

0.08

EPR EPR

13– 238 S 26– 288 and 358 S 57– 638 N

2,600– 2,900 2,450

7.7 8

2

10 6

600– 1,825

1

7

(7)

0.08

0.08

Peate et al., 2001

37– 418 N

950– 3,000

1

23

0.13

Bourdon et al., 1996a,b, 2000

MAR (FAMOUS) MAR MAR

368 500 N 29– 308 N 338 S

1.31 (1.47)

0.55

Condomines et al., 1981 Bourdon et al., 1996b Lundstrom et al., 1998a

AAD

508 S

2,550 3,300– 4,000 ,3,500 and 2,700 ,4,200

MAR (Reykjanes Ridge) MAR (FAZAR)

1.1 1.2 1.8

3 2 6

3.7

2

20

1

6

1

(6)

0.44 (0.46)

Bourdon et al., 1996b

The criteria for selection of samples from full U-series database are described in Section 3.14.3.1. The numbers of suitable samples available for each U-series disequilibrium are listed, and those determined by alpha counting indicated by italics. The equiline gradient is the regressed slope of sample arrays on an equiline diagram (see Figure 7), together with a standard error on the slope. Arrays for locations with less than six samples were not calculated. In some cases two slopes are reported, which include a preferred slope and a regression of all data (in parentheses). Preferred slopes exclude seamount data in JDF and 8– 108 N EPR arrays, which potentially have anomalously high melting rates and naturally lower disequilibrium than normal ridge samples. The preferred fit for 8 – 108 N EPR only includes the 98 300 – 98 500 N axial sample set. The preferred fit in the FAZAR samples includes only the samples from the most robust segment (KP4) which are likely best located and have experienced least post-eruptive decay. However, in all cases, the preferred and inclusive regressions are close to error of each other.

474

Melt Migration in Oceanic Crustal Production 3.0 Age uncertainty for axial samples from: Fast spreading ridge

2.8

Slow spreading ridge 2.6 2.4

Activity

2.2 2.0 ( 230 Th / 238 U )

1.8

( 226 Ra / 230 Th) ( 231Pa / 235 U )

1.6 1.4 1.2 1.0 0

50,000

100,000

150,000

200,000

Time (yr)

Figure 4 Decay of typical initial U-series excesses in MORB. Indication of age uncertainty for samples located within the axial neovolcanic zone for slow and fast spreading ridges indicated by double headed arrows (see text for details).

is significant for (231Pa/235U), largely because of its greater initial excess relative to measurement uncertainty. From Figure 4 it is clear that reliable data on “initial” (226Ra/230Th) can only be obtained on samples younger than a few hundred years, whereas samples younger than ,10 kyr provide good estimates on initial (230Th/238U) and reasonable values of initial (231Pa/235U). It is, thus, necessary to have good independent controls on the age of samples. In the sub-aerial environment, U-series analyses are generally made on samples whose ages are documented by observations and historical records. In some cases, older samples can be accurately corrected from 14 C ages of underlying charcoal, for example. Such constraints are rarely available in studying MORBs, although there are now a few samples from recent eruptions as a result of concerted monitoring programs (Chadwick et al., 1991, 1998) and serendipity (Haymon et al., 1993). There are some qualitative indices of sample age, e.g., state of glass preservation, sediment cover, thickness of Mn crust (Ballard et al., 1981; Haymon et al., 1991; Macdonald et al., 1988). Calibration of such observations with U-series methods is sparse and one study has pointed out inconsistencies between the two approaches (Sturm et al., 2000). Important age constraints are available, however, from the limited width of the neovolcanic

zone on mid-ocean ridges. Assuming lavas are erupted randomly within the axial rift, a measure of the uncertainty in their age can be gauged from the width of the neovolcanic zone divided by the spreading rate (determined over a longer time period from magnetic anomalies). Taking typical neovolcanic zones (see Perfit and Chadwick, 1998) for fast spreading ridges (200 m wide, full spreading rate 80 mm yr21) and slow spreading ridges (2 km wide, full spreading rate 20 mm yr21), this approach gives age uncertainties ranging from 2.5 kyr at fast spreading ridges to 100 kyr at slow spreading ridges. Thus for fast-spreading ridges, all axial samples are likely to give a robust measure of eruptive ( 230Th/ 238U) and ( 231Pa/235U), but not (226Ra/230Th). For slow-spreading ridges, even samples well located within the axial rift zone potentially require age correction for all disequilibria. Clearly, this is an oversimplified model that does not consider manifest complexity in crustal accretion (e.g., Perfit and Chadwick, 1998; Sims et al., in press; White et al., 2002), but does highlight the basic issue. It is also worth stressing that good sample location and associated bathymetry is critical to the effectiveness of these basic age constraints. Locations of samples taken by submersible are much more precise than from those from dredging that have a spatial resolution of ,300 –500 m at best.

Observations 226

230

Measured ( Ra/ Th) even on the fastestspreading ridges can thus only be taken as minimum estimates of primary values. Observations at New Mounds site at ,458 N on the Juan de Fuca (JDF) Ridge (Chadwick et al., 1991), at ,428 400 N on the North Gorda Ridge (Chadwick et al., 1998) and BBQ site at ,98 500 N on the northern East Pacific Rise (EPR) (Haymon et al., 1993), suggest eruptive activity in the last decade and so samples from these sites should yield reliable initial (226Ra/230Th). 210Po – 210Pb dating (t1/2 210Po 134 d) has verified the young age of the BBQ and N. Gorda site samples (Rubin et al., 1994, 1998). This technique could be used elsewhere in combination with morphological observations to identify other “zero age” samples. It also might be possible to use 210 Pb – 226 Ra disequilibrium (t1/2 210Pb , 25 yr) to identify additional samples that are young enough to preserve primary (226Ra/230Th). 210Pb – 226Ra disequilibrium is observed in some sub-aerial basalts (Gauthier and Condomines, 1999; Sigmarsson, 1996), and like 210 Po – 210Pb disequilibrium can be related to preferential loss of a more volatile nuclide during shallow degassing (Gauthier and Condomines, 1999). A promising initial study of 210Pb – 226Ra disequilibrium in MORB was reported by Rubin et al. (2001). An additional uncertainty in interpreting the (226Ra/230Th) data is the time a magma batch resided in the oceanic crust, prior to eruption. Cooper et al. (2003) have analyzed plagioclase separates from the historic N. Gorda flow to place constraints of ,1.5 ka on the crustal residence time of magmas before eruption. Nevertheless, it would be useful to have yet tighter constraints on the residence time to define better initial (226Ra/230Th). It is possible that shorter-lived, mantle-derived disequilibria (e.g., 231Pa – 228Ac) could provide such information. For the longer-lived disequilibria, crustal residence time is likely unimportant and age constraints less critical. Crucially, the presence of any (226Ra/230Th) disequilibrium implies samples have near initial (230Th/238U) and (231Pa/235U), Figure 3. (226Ra/230Th) disequilibrium is thus an especially valuable tool for screening (230Th/238U) and (231Pa/235U) on intermediate and slow spreading ridges. In compiling our database, we have included well located samples from within 1 km of the axis of intermediate or fast-spreading ridges, and any samples that show significant (.3%) 226Ra-excesses. The latter constraint has been used to exclude a number of samples from slow-spreading ridges. We have also included two samples sets from slow-spreading ridges for which there are no 226Ra data (Bourdon et al., 1996b). These suites are notable in one case for high (230Th/238U) and in the other for a modest

475

230

Th-deficit, neither of which can be attributed to post-eruptive decay.

3.14.3.3 General Signatures of Disequilibrium Data and interest in disequilibrium in MORB historically started with (230Th/238U), followed by (226Ra/230Th) and finally (231Pa/235U). We follow a similar path in reviewing the general magnitude of disequilibria in MORB (Figure 5). Subsequently we explore finer scale structure and variability of disequilibria with other geochemical and geophysical parameters (Section 3.14.3.5). Before considering disequilibrium, however, we examine (234U/238U), that should not be fractionated by high temperature processes (Section 3.14.2.1). Figures 5(a) and (b) illustrate that this is indeed the case for our MORB compilation. Both mass-spectrometric and alphacounted data sets are dispersed around unity by an amount consistent with their respective typical analytical uncertainty, although some individual studies provide considerably more precise data (e.g., Sims et al., 2002). It is now well established that MORBs typically show modest 230Th-excesses of 10 –25% (Figures 5(c) and (d)). This signature, evident in the first U-series measurements of MORBs from the FAMOUS region of the Mid-Atlantic Ridge (MAR), was initially a surprise (Condomines et al., 1981). Subsequent work verified significant disequilibrium in MORBs (Ben Othman and Alle`gre, 1990; Newman et al., 1983; Reinitz and Turekian, 1989; Rubin and Macdougall, 1988). As discussed above, alpha-counting techniques were challenged by the low uranium concentrations of MORBs. Typical precisions (5 – 15%) were often close to the overall degree of disequilibrium and the case for primary disequilibrium was debated (Krishnaswami et al., 1984). Mass-spectrometric (230Th/238U) data typically have a precision of ,1– 2%, which provides a finer resolution picture of disequilibrium. For example, there is only a single mass-spectrometrically analyzed sample with a 230Th-deficit, and very few have 230Thexcesses less than 5%. Yet the overall picture provided by both alpha and mass-spectrometric techniques is reassuringly similar. 226 Ra– 230Th disequilibrium in MORBs is much more dramatic than 230Th – 238U disequilibrium, with (226Ra/230Th) as great as 4.2 (Figures 5(d) and (e)). Alpha-counted and mass-spectrometric data both document (226Ra/230Th) ratios in excess of 2. Alpha-counted errors on (226Ra/230Th) are ,10% (Rubin and Macdougall, 1988) and larger than the ,1% errors by mass-spectrometry (e.g., Volpe and Goldstein, 1993). In contrast to (230Th/238U) disequilibrium, the difference in

Melt Migration in Oceanic Crustal Production 45

9

40

8

35

7

30

6 Frequency

Frequency

476

25 20

5 4

15

3

10

2

5

1

0 0.98

0 0.99

1.00

(a)

1.01

1.02

1.03

0.90

1.00

(b)

234U/238U)

( 20

1.10

(234U/238U)

15

Frequency

Frequency

15

10

10

5 5

0 0.8

0 0.9 (c)

1

1.1

1.2

1.3

1.4

1.5

(230Th/238U)

(d)

1

1.2

1.4

1.6

1.8

(230Th/238U)

Figure 5 Histograms of activity ratios for key U-series nuclide pairs in compiled MORB data set: (a) (234U/238U) by mass spectrometry, using reported (234U/238U). In the absence of citied l234 for all studies we did not try correcting to a common l234, (b) (234U/238U) by alpha counting, note expanded scale relative to (a), (c) (230Th/238U) by mass spectrometry, (d) (230Th/238U) by alpha counting, (e) (226Ra/230Th) by mass spectrometry in all samples from intermediate to fast spreading ridges (dark gray) and definitive historic events (light gray), (f) (226Ra/230Th) by alpha counting, (g) (231Pa/235U) in all samples (dark gray) and those from definitively historic eruptions (light gray).

the precision of the techniques is not highly significant given the much larger degrees of 226Ra – 230Th disequilibrium. Much more important are age uncertainties, and as discussed above (Section 3.14.3.2) most of these 226Ra data are minimum values. Only the historic samples highlighted (Figure 5(e)) are clearly reliable indicators of initial (226Ra/230Th). Interestingly, the historic samples do not show strikingly higher (226Ra/230Th) than a sizable proportion of the mass-spectrometry data set (Figure 5(e)). This suggests that many of the lavas sampled are under a few hundred years old. 235 U is 137.88 times less abundant than 238U, and so (231Pa/235U) measurements are analytically very challenging. Only a small number of massspectrometric measurements have been reported, but the results are striking. 231Pa-excesses are

large (Figure 5(g)), although not quite as extreme as some (226Ra/230Th) ratios (Figure 5(f)). The current data set shows small variations of (231Pa/235U) around a rather well-defined median of ,2.6. Historic samples are also shown and reassuringly the less precisely dated samples do not appear distinct from the historic eruptions (Figure 5(g)). The presence of significant 230Th – 238U and large 226Ra – 230Th and 231Pa – 235U disequilibria in MORBs is a first-order observation. Accounting for the magnitude of these disequilibria places valuable constraints on melt migration and required development of a new series of melting models (Section 3.14.4.3). However, before ascribing disequilibria to the melting process, it is first necessary to discuss the possible role of crustal contamination.

Observations

477

7

10 all intermediate/fast historic

8

6

Frequency

Frequency

5 6

4

4 3 2

2 1 0

0 1

(e)

2

3

4

5

1

(226Ra/230Th)

2

3

4

5

(226Ra/230Th)

(f)

12 all data

10

historic

Frequency

8 6 4 2 0 1 (g)

1.5

2

2.5

3

3.5

4

(231Pa/235U)

Figure 5 (continued).

3.14.3.4 Assessment of Shallow-level Contamination Assimilation of near-surface material that has low-temperature-induced disequilibrium is a spectre frequently raised in the discussion of U-series disequilibria in MORBs. For example, ferromanganese coatings and hydrothermal precipitates on seafloor samples sequester significant amounts of unsupported 234U, 230Th, and 231Pa from seawater (German et al., 1993; Lalou et al., 1993; Shimmield and Price, 1988). Inclusion of even small amounts of such material in the measured lavas can induce significant disequilibrium (Figure 6). To minimize this potential problem, analyses are nearly always made on carefully handpicked glasses. Picking the large amounts of material required for older alphacounting measurements was an onerous task, and it was therefore difficult to screen rigorously for contaminant on such large samples. Newman et al. (1983) noted that some of their samples had distinctly higher (230Th/232Th) and (230Th/238U)

than others. They attributed this to contamination and did not further consider these analyses. Similarly, 234U and 230Th contamination of some samples was implicated in a study of mid-Atlantic MORB samples by Condomines et al. (1981). In addition to hand picking glass, various leaching procedures have been developed (e.g., Goldstein et al., 1989; Reinitz and Turekian, 1989) in order to remove any ferro-manganese coating not spotted during picking (e.g., in small cracks). It is clearly important that such preparation techniques do not themselves cause fractionation of U-series nuclides in leached glass. Contamination may also occur before eruption, by magmatic assimilation of altered material. Chemical proxies are required to assess such contamination. Bourdon et al. (2000) have recently comprehensively reassessed the issue of contamination using 10Be. The presence of any of this short-lived (t1/2 , 1.5 Myr), cosmogenically produced nuclide in MORBs should be attributable to near-surface contamination alone. Using typical values of 10Be and U-series nuclide

478

Melt Migration in Oceanic Crustal Production 4.00

0.005

3.50

(231 Pa/235 U)

3.00

2.50

0.002

2.00 234U excess less than

analytical uncertainty

1.50

1.00 1.00

0.001 0.0005 0.0002 1.10

1.20

1.30

1.40

1.50

1.60

(230 Th/238 U)

Figure 6 Illustration of the effects of contamination. (230Th/238U) and (231Pa/235U) for compiled MORB data set plotted together with a mixing line for different weight fractions (labeled) for ferro-managanese coating added to a putative MORB with no primary U-series disequilibria. The contaminant calculation is after Bourdon et al. (2000) with 50 ng g21 [U] MORB in secular equilibrium and 1 mg g21 [U] ferro-managanese coating with (234U/238U) ¼ 1.15, (230Th/238U) ¼ 6.25, (231Pa/235U) ¼ 30. The shaded box indicates the maximum values that could be ascribed to contamination alone for the ,98 500 N East Pacific Rise data set of (Sims et al., 2002), where equilibrium (234U/238U) values have been measured with an average precision of 2‰.

concentrations for ferro-manganese materials, a maximum of between 15% and 1% of the 231Pa and 230Th-excesses, respectively, can be accounted for by contamination (Bourdon et al., 2000). The study also highlighted the effectiveness of hand picking in producing reliable samples. (234U/238U) ratios in MORBs are also perturbed by addition of ferro-manganese material. Unfortunately, (234U/238U) is not routinely measured with a suitable precision to clearly rule out contamination. The recent work of Sims et al. (2002), however, document high precision (234U/238U) measurements of EPR samples that are generally within 3‰ of equilibrium. Figure 6 illustrates the maximum effect this uncertainty implies for possible contamination, and shows that the maximum permissible contamination cannot account for the high measured (231Pa/235U) in particular. Sims et al. (2002) also investigated other sensitive geochemical tracers of seawater related contamination, d11B and C1/K ratios, and found no evidence that these were elevated or correlated with disequilibria. The issue of shallow-level processes in generating 226Ra – 230Th disequilibrium is somewhat different to that for 238U – 234U – 230Th and 235U – 231Pa. Ra is not strongly particle reactive and not strongly concentrated in

ferro-manganese coatings. Thus to invoke contamination to explain 226Ra-excesses at the same time as 230Th- and 231Pa-excesses requires two unrelated processes. Hydrothermal processes at the ridge axis result in precipitation of barite particles onto the ocean floor, which initially have extreme 226Raexcesses (e.g., Moore and Stakes, 1990). Minor addition of recent hydrothermal barite into an MORB flow would dramatically increase any (226Ra/230Th). Although this seems an attractive mechanism to produce what were initially puzzlingly large 226Ra-excesses in MORB, there is a strong chemical argument against this explanation. The effect of barite contamination should also be evident in the barium contents of affected MORB, but not on other highly incompatible elements such as Rb. Unless the amount of barite contamination is constant worldwide, in locations with and without active hydrothermal venting, then it becomes very difficult to account why Ba/Rb remains near constant over a global MORB data set (Hofmann and White, 1983). It becomes even more difficult to account for the near identical ratio of Ba/Rb in MORB and OIB (Hofmann and White, 1983), as the latter are erupted in a setting where barite production is significantly lower. It is also worth noting that

Observations 226

some of the highest Ra-excesses occur in the most incompatible element depleted samples. One would expect the addition of barite barium to be prominent in these samples. It has recently been suggested that plagioclase accumulation can account for 226Ra-excesses (Saal et al., 2002a,b; Van Orman et al., 2002b). Such a mechanism seems unlikely to explain the general presence of 226Ra-excesses in recent MORBs. The partition coefficient of radium in plagioclase is greater than thorium but is nevertheless less than 1 (Blundy and Wood, 2003b). Thus, the dominant reservoir of radium will be the melt and a large amount of plagioclase contamination is required to influence a given amount of magma (Cooper et al., 2003). In any individual case, the role of plagioclase contamination can be further explored using other geochemical tracers, e.g., Sr (Saal et al., 2002a,b). Given the arguments, most observed disequilibria can be ascribed with some confidence to primary mantle processes. They, thus, form first order tests of any model of melting and melt migration. 3.14.3.5 Further Observations In addition to the basic observations of U-series excesses (Section 3.14.3.3), further insight into processes in the melting regime is gained by investigating how the disequilibria co-vary with each other and additional geochemical and geophysical parameters. The quality of correlations is assessed using a Spearman rank correlation coefficient, rs, and a probability factor, p. rs ranges from 21 to 1; 21 is a perfect negative correlation, 0 is no correlation, and 1 is a perfect positive correlation. The rank correlation is more robust than the standard linear correlation as it is insensitive to outliers (see Press et al., 1992). Moreover, it is possible to calculate a probability p that a larger values of rs could not be produced at random (i.e., large values of p imply high confidence that the correlation could not be better for a random distribution). The systematics discussed below help distinguish between a range of different melting models (see Section 3.14.4) that can explain the basic magnitudes of observed disequilibria (see Section 3.14.3.3), but have rather different implications for melt migration.

by chemical fractionation (on timescales short compared to the half-life of 238U and 232Th), while the y-axis (230Th/232Th) is sensitive to the decay or production of 230Th and will change on the 230Th timescale (Figure 7(a)). Samples in secular equilibrium will all plot on the equiline which has a slope of unity, i.e., (230Th) ¼ (238U); therefore, (230Th/232Th) ¼ (238U/232Th). Contours of constant disequilibrium are represented by straight lines that fan out from the origin (Figure 7(a)). Variable fractionation of uranium and thorium generates a horizontal array of samples (Figure 7(a)). If the melt and residue remain in chemical equilibrium, this flat array can persist indefinitely (Section 3.14.2.2.2). However, if the melts and residues become chemically separated, the array rotates with age, pivoting around a point on the equiline, until it ultimately returns to equilibrium with a slope of 1. As in any isochron plot, a linear array may have age significance, or it may simply reflect mixing between two unrelated samples. The x-axis of the equiline diagram is an inverted form of the Th/U ratio, which is frequently used as an index of “enrichment.” Enriched samples, with higher incompatible element contents tend to have higher Th/U, or lower (238U/232Th). The overall trend can be observed in our MORB data set (Figure 3). Figure 7(b) shows the global MORB data set plotted on an equiline diagram. The global array cuts across lines of equal – disequilibrium such that “enriched” samples with low (238U/232Th) tend to have high degrees of 230Th – 238U disequilibrium and “depleted” samples with high (238U/232Th) generally have lower degrees of disequilibrium. “Local” data sets from limited geographical regions (Figures 7(c) – (f)) show similar arrays: samples with the lowest (238U/232Th) display the greatest degrees of disequilibrium (Goldstein et al., 1991). In some instances, local arrays define a significant portion of the total global array (e.g. Figures 7(c) –(f) and Lundstrom et al., 1999). The higher precision of mass-spectrometric measurements has identified rather linear sample arrays in the equiline plots for some locations (Lundstrom et al., 1998b). The best-fit slopes of these arrays are variable (Lundstrom et al., 1998b) and values for the better defined arrays in our data set are reported in Table 1. 3.14.3.5.2

3.14.3.5.1 230

Equiline diagram

Th – 238U disequilibrium data is frequently plotted on the so-called equiline diagram (Alle`gre, 1968). This plot displays 230Th– 238U disequilibria in a manner analogous to a traditional isochron diagram. The x-axis (238U/232Th) is affected only

479

Source enrichment

In both the local and global arrays on the equiline, there is a general trend from low degrees of disequilibrium at high (238U/232Th), or (230Th/232Th), to higher degrees of disequilibrium at low (238U/232Th), see Figure 7(b). Decreasing (238U/232Th) ratios are associated

480

Melt Migration in Oceanic Crustal Production 1.6

ss

ine

ce

0 Th 23

uil

ex

Eq

%

20

(230Th/232Th)

1.4

1.2

cit

efi

hd 30 T

Ingrowth

2

%

20

Melt

Elemental fractionation

Residue

Decay

1.0

0.8 0.8

1.0

1.2

1.4

1.6

238U/232Th)

(

(a) 1.6

(230Th/232Th)

1.4

1.2

1.0 all data 8–10˚ N EPR

0.8 0.8 (b)

1.0

1.2

1.4

1.6

(238U/232Th)

Figure 7 Equiline plots. (a) Cartoon of systematics of equiline diagram, (b) all compiled MORB data, open symbols highlight samples from 8 – 108 N EPR to indicate substantial fraction of global array evident in a single geographically restricted sample set, (c) samples from Gorda and Juan de Fuca ridges, (d) samples from 128 and 218 N EPR, (e) samples from 8 –108 N EPR, including Siqueiros transform, (f) samples from Azores platform (FAZAR), shown by segment and Reykjanes Ridge. Data sources can be found in Table 1. Typical precision for massspectrometric measurement comparable to sample size and for clarity errors are only shown when significantly larger. Cited counting errors (2s) shown on alpha counting measurements. Short dashed and long dashed lines indicate 230 Th excesses of 20% and 10% respectively.

Observations

481

1.6

(230Th/232Th)

1.4

1.2 Juan de Fuca Axial seamount (JdF) Gorda

1.0

0.8 0.8

1.2

1.0

1.4

1.6

238U/232Th)

(

(c) 1.6

(230Th/232Th)

1.4

1.2

21˚ N EPR 12˚ 50′ N EPR

1.0

0.8 0.8

1.0

1.2

(d)

1.4

1.6

238U/232Th)

(

Figure 7 (continued).

with “enrichment” of incompatible element abundances and ratios of more to less incompatible element ratios (e.g., Figure 3). There is, thus, some evidence for the co-variation of the degree of (230Th/238U) disequilibrium with “enrichment.” A key question is what caused the enrichment and when did it occur? (238U/232Th) ought to be a rather robust index of source enrichment. As highly incompatible elements (Section 3.14.4.1), uranium and thorium

are likely to be fractionated by the melting process only at very low degrees of melting (Section 3.14.2.2). Average MORB is thought to be the aggregate of ,10% melting (Klein and Langmuir, 1987; McKenzie and Bickle, 1988), in which there should be no net elemental fractionation of uranium from thorium. Whilst average MORB is the product of a large degree of melting, individual MORB samples likely represent an incomplete blend of melts produced at different

482

Melt Migration in Oceanic Crustal Production 1.6

(230Th/232Th)

1.4

1.2 EPR 9–10˚ N Siqueiros transform Lamont Seamount

1.0

0.8 0.8

1.0

1.2

(e)

1.4

1.6

238U/232Th)

(

1.6

(230Th/232Th)

1.4

FAZAR KP1 (FAZAR) KP2 (FAZAR) KP4 (FAZAR) Reykjanes Ridge

1.2

1.0

0.8 0.8

1.0

1.2

(f)

1.4

1.6

(238U/232Th)

Figure 7 (continued).

locations in the melting regime (Langmuir et al., 1977; Sobolev and Shimizu, 1993). The presence of 230Th – 238U disequilibrium indeed testifies to the fractionation of thorium from uranium some where during the melting process, even if net elemental fractionation should not be evident in a true mean of all the melts produced (Section 3.14.4). In order to distinguish between source heterogeneity and recent melt induced fractionation, it

would be useful to have tracers that are not fractionated by “melting process.” Long-lived radiogenic isotope ratios are such tracers. Sadly, long-lived radiogenic isotope data are rarely available on the same samples analyzed for (230Th/238U). Since (238U/232Th) is available for every sample with measured (230Th/238U), this imperfect trace element ratio is most commonly used as a proxy for source composition.

Observations The few data sets with long-lived radiogenic isotopes as well as 230Th – 238U data can be used to assess how closely (238U/232Th) variations correlate with more robust tracers of heterogeneity (Figure 8). This is hence not a comprehensive plot of all (238U/232Th) and 87Sr/86Sr data for MORBs, but only for those samples in our data set for which the systematics of disequilibria are also well constrained. A general anticorrelation of 87 Sr/86Sr with (238U/232Th) is evident, which represents the so-called Sr – Th array (Condomines et al., 1981; Rubin and Macdougall, 1992). The samples in Figure 8 comprise subsets from six main locations. Thus, in the same way that variations of slopes on equiline were examined on a “local” scale in Figure 7, it is important to assess correlations on (238U/232Th) with 87Sr/86Sr on a similar scale. The FAZAR samples from the Azores platform (Bourdon et al., 1996a) show a rather significant isotopic gradient along the sample section, but variations of (238U/232Th) that occur on a segment length are comparable to those across the whole geochemical anomaly. The Reykjanes Ridge samples analyzed for U-series nuclides (Peate et al., 2001) are taken from a region of limited isotopic variability (Murton et al., 2002) on the general trend from “normal” Atlantic MORB to enriched Icelandic compositions (Schilling, 1973). Despite the minor 87Sr/86Sr variation in

483

the Reykjanes Ridge samples, there is significant, unrelated (238U/232Th) variability. The suite from 128 500 N EPR (Ben Othman and Alle`gre, 1990) shows the clearest covariation of (238U/232Th) with 87Sr/86Sr and indeed was used as a case example of the marble-cake mantle (Prinzhoffer et al., 1989). In contrast, samples from , 98 300 – 98 500 N EPR (Goldstein et al., 1993; Sims et al., 2002) show much lower (238U/232Th) than nearby depleted MORB from the Siqueiros transform (Lundstrom et al., 1999) but all these samples have comparable 87Sr/86Sr (Sims et al., 2002). We note, however, that one E-MORB from the Siqueiros transform shows both low (238U/232Th) and elevated 87Sr/86Sr (Lundstrom et al., 1999; Sims et al., 2002). This sample is not plotted on Figure 8 as it does not meet the age criteria of our compilation, and so interpretation of its disequilibria is uncertain. Samples from both the JDF and Gorda ridges are striking in the lack of correlation between (238U/232Th) and 87Sr/86Sr (Goldstein et al., 1991; Sims et al., 1995). Yet, the sample suites from these localities define some of the better arrays on equiline plots (Figure 7(c)). Notably, the most enriched sample from the JDF, with lowest (238U/232Th) and highest (230Th/238U) has a markedly unradiogenic 87Sr/86Sr (Figure 8). A clear link between (238U/232Th) and ancient source enrichment is thus not readily evident on

1.6 Siqueiros depleted 8 –10˚ N EPR

1.5

depleted Siqueiros (8˚ 20′ N) 12˚ 50′ N EPR Gorda Juan de Fuca (JdF)

1.4

Reykjanes Ridge

(238U/232Th)

1.3

JdF/Gorda depleted

Azores Platform

1.2 1.1 1.0 0.9 0.8 0.7022

JdF/Gorda enriched

0.7024

0.7026

0.7028

0.703

0.7032

0.7034

0.7036

(87Sr/86Sr)

Figure 8 87Sr/86Sr versus (238U/232Th) for the MORB compilation of this study. Hence only samples with accompanying U-series data, not effected by post-eruptive decay, are plotted. Despite an overall trend of decreasing (238U/232Th) with increasing 87Sr/86Sr, on a local scale, samples span a wide range of (238U/232Th) at similar 87 Sr/86Sr. Additional sources for 87Sr/86Sr data: (Davis and Clague, 1987; Dosso et al., 1999; Murton et al., 2002; Sims et al., 1995).

484

Melt Migration in Oceanic Crustal Production

a local scale. Whilst a few enriched samples have elevated 87Sr/86Sr, more frequently this is not the case. Moreover, there are large ranges of (238U/232Th) in samples with similar 87Sr/86Sr. Since Lundstrom et al. (1998b) noted that samples with lower (238U/232Th) had higher (230Th/238U) (i.e., plotted further to the left of the equiline) and inferred (238U/232Th) variations to be related to source heterogeneity, they required a mechanism to produce larger degrees of disequilibrium when melting a more enriched source. Lundstrom et al. (1995) accounted for this by invoking small-scale, enriched heterogeneities with a significantly different lithology (garnet–pyroxenite) to the surrounding peridotite (Alle`gre and Turcotte, 1986), and later expanded on this model (Lundstrom et al., 2000). In the light of Figure 8, however, additional processes are likely active. The possible fractionation of (238U/232Th) as a result of MORB generation is discussed in Section 3.14.4. 3.14.3.5.3

Variations of disequilibria with spreading rate

The local arrays of data on equiline plots show variations in slope from region to region (Figure 7), with a much steeper trend for samples from the 8– 108 N EPR for example (Figure 7(e)) than the near horizontal array for the Reykjanes Ridge (Figure 7(f)). This was originally noted by Lundstrom et al. (1998b), who suggested a relationship between spreading rate and such local equiline slopes. Slow-spreading ridges apparently define flatter arrays than fast-spreading ridge segments. Using our data compilation, we fitted straight lines to the local arrays illustrated in Figure 7. The results are reported in Table 1 and plotted against spreading rate in Figure 9(a). Correlation of local equiline slopes with spreading rate is very poor (rs ¼ 20.25, p ¼ 0.4), unless the data set from the Azores platform (FAZAR samples) and other less well defined arrays are excluded. One key data set used by Lundstrom et al. (1998b), from the ultrafast spreading S. EPR (Rubin and Macdougall, 1998), is not plotted here. These data define the steepest normal ridge segment array, lying close to the equiline. However, all bar two of these samples have been excluded from this study due to potential problems of blank contamination (Section 3.14.3.1), that result in very high (230Th/232Th) and (238U/232Th). Given the doubts raised over such extreme values above, we have not included the S. EPR samples in this compilation. The recent repeat analyses of some of these samples by mass-spectrometry (Rubin et al., 2000) will, thus, be of significant interest. Our filtering of

some alpha-counted data has also resulted in a different slope for our compiled ,128 N samples compared to the value of (Lundstrom et al., 1998b), 0.05 versus 0.16, respectively. The Reykjanes Ridge and Azores platform regions represent the slow-spreading end-members of the global data set. Both of these locations are anomalously shallow and inferred to be affected by mantle that feeds nearby ocean islands (Iceland and the Azores, respectively). Despite these similarities in their geophysical setting, the systematics of samples from these two locations are strikingly different (Figure 7(f)). The Reykjanes data pin the Lundstrom et al. (1998b) global trend, whereas the FAZAR data greatly degrade any correlation (Figure 9(a)). It is not clear why the FAZAR samples should be different if the Lundstrom et al. (1998b) model is globally applicable. Whilst the FAZAR samples do cover a wider range of axial depths than other sample suites shown in Figure 9 (see implications of this in Section 3.14.4), individual ridge segments span a much smaller range of axial depth and display similar slopes to the overall trend (Figures 7(f) and 9(a)). It is also worth noting, however, that it is the slope of arrays on an equiline diagram that Lundstrom et al. (1998b) related to spreading rate and not absolute 230Th-excesses. Lundstrom et al. (1998b) pointed out that the slope of the equiline represented one way of characterizing a range of compositions from a single location. Interestingly, a simple correlation of (230Th /238U) with spreading rate is more significant (rs ¼ 20.57, p ¼ 1) than for the equiline slopes. (231Pa / 235U) variations are uncorrelated (rs ¼ 20.16, p ¼ 0.68) with spreading rate (Figures 9(b) and (c)). The degree of disequilibrium produced during “in-growth” models is inversely dependent on melting rate (see Section 3.14.4), which in the MORB can be linked to upwelling rate. It is, thus, perhaps encouraging to see some apparent inverse spreading rate dependence in the disequilibria data (Figure 9(b)). However, any relationship between spreading rate and erupted disequilibria is not likely to be straightforward (see Section 3.14.4.3.3). It is thus not necessarily surprising that (230Th/238U) shows a weak correlation with spreading rate but (231Pa/235U) does not. It may be significant that the slowest-spreading ridges are also the shallowest (see next section), and that several effects may combine to produce correlations observed. 3.14.3.5.4

Variations of disequilibria with axial depth

Bourdon et al. (1996b) presented a data set that indicated a general correlation of segment averaged (230Th/238U) and axial depth. Key data

Observations

485

2 alpha counted preferred model alternative fit

Azores platform (FAZAR)

Equiline slope

1.5

1 8–10˚ N EPR 23˚ S MAR

Juan de Fuca

0.5 Gorda 12˚ N EPR

Reykjanes ridge

0

21˚ N EPR

–0.5 0

1

2

(a)

3

4

5

6

Half-spreading rate (cm yr –1) 1.8

(230Th/238U)

1.6

1.4

1.2

1.0

Equilibrium

0.8 0

(b)

1

2

3

4

5

6

7

8

9

Half-spreading rate (cm yr –1)

Figure 9 (a) Equiline slope versus half-spreading rate for localities plotted in Figure 7, (b) (230Th/238U) versus half spreading rate for all compiled MORB samples, (c) (231Pa/235U) versus half spreading rate for compiled MORB samples. Data for the regressed equiline slopes and half spreading rates are reported in Table 1. Where reported in Table 1, preferred and inclusive regression are plotted.

on this plot were the shallow FAZAR samples with high degrees of (230Th/238U) and samples from two deep ridge settings, the Australian Antarctic Discordance (AAD) and 308 N MAR, with (230Th/238U) close and even less than unity. Unfortunately none of these samples have

measured (226Ra/230Th) excesses to assess posteruptive decay. Yet, as well located samples from an intermediate spreading center, it is unlikely that uncertainty in age accounts for the very low disequilibrium in the AAD samples. Moreover, the 238U-excess in a single sample

486

Melt Migration in Oceanic Crustal Production 4.0

3.5

(231Pa/235U)

3.0

2.5

2.0

1.5

1.0 0 (c)

1

2

3

4

5

Half-spreading rate (cm

6

7

8

9

yr –1)

Figure 9 (continued).

from 308 N MAR cannot simply be the result of post-eruptive decay. The FAZAR samples have some of the highest (230Th/238U) observed on the ridge system. Any post-eruptive decay in these samples should only lower their disequilibrium and so not grossly influence the global correlation. Bourdon et al. (1996b) accounted for their correlation as the result of a deeper onset of melting of hotter (or more enriched) mantle. Deeper melting regimes have larger amounts of melting in the garnet stability field, with higher DU/DTh and this can result in higher (230Th/238U) of the erupted melts (see Section 3.14.4). Melting regimes beneath the deepest ridges potentially do not extend into the garnet stability field and so modest 230Th-excesses or even deficits could be expected (again see Section 3.14.4). Since the publication on Bourdon et al. (1996b), the understanding of uranium and thorium partitioning in clinopyroxene has improved considerably. A continuous increase in DU/DTh in clinopyroxene with increasing pressure provides a further mechanism to account for the global correlation (e.g. Landwehr et al., 2001). Figure 10 shows an updated plot of (230Th/238U) against axial depth, that includes data from the shallow Reykjanes Ridge (Peate et al., 2001) and deep southern MAR (Lundstrom et al., 1998a) published since Bourdon et al. (1996b). A correlation persists (rs ¼ 20.58, p ¼ 1), although the new Reykjanes Ridge data do not clearly substantiate the original trend. A single Reykjanes

sample shows very high (230Th/238U), but this is not true of the whole data set. Moreover, recent unpublished data (not shown) from the shallow Kolbeinsey Ridge (Sims et al., 2001) is also not generally supportive of the Bourdon et al. (1996b) correlation. Kolbeinsey samples define a near horizontal slope on the equiline diagram. While the highest degree of disequilibrium is quite high (230Th-excesses of 24%), the Kolbeinsey ridge also boasts samples with 238U-excesses (Sims et al., 2001). More data from the Kolbeinsey Ridge are needed to see if the 238U-excesses reported by Sims et al. (2001) significantly influence the segment average. As discussed above (Section 3.14.3.5.3) 230Thexcesses show a negative correlation with spreading rate of similar significance to the negative correlation with axial depth. Moreover, the samples that define the shallowest ridges are from the slowest spreading ridges. In order to disentangle these effects it would be useful to have more data from locations such as the southern EPR, which is fast but not anomalously deep, and the (AAD) which is deep but does not have an extreme spreading rate. A key point about Figure 10 (and Figure 9(b)) is that (230Th/238U) variability at a given depth (or spreading rate) can be a major proportion of the total global range. This further emphasizes the importance of local variations discussed in the previous section. Whilst the global trend is still significant despite this scatter, it is clear that local variations in the melting process have as

Observations

487

1.8 all samples

1.7

FAZAR samples

1.6

(230Th/238U)

1.5 1.4 1.3 1.2 1.1 1.0 0.9 0.8 0

1,000

2,000

3,000

4,000

5,000

Axial depth (m)

Figure 10 (230Th/238U) against axial depth for all compiled MORB samples. Horizontal dashes further indicate FAZAR samples which cover a significant portion of the array (Bourdon et al., 1996a), but are not the sole source of the correlation.

important an effect on 230Th – 238U disequilibrium as the ,200 8C difference in mantle temperature invoked to account for changes in axial depth (Bourdon et al., 1996a; Klein and Langmuir, 1987). This observation has resonances with local variations in the major element systematics of MORB (Klein and Langmuir, 1989; Langmuir et al., 1992).

3.14.3.5.5

Mutual covariations of disequilibria

So far, we have only considered 230Th – 238U disequilibrium, for which there is the largest body of data, least effected by post-eruptive decay. It is now worth exploring the mutual variations of the different disequilibria. As discussed above (Section 3.14.3.2) and illustrated in Figure 4, it is difficult to obtain robust primary estimates of (226Ra/230Th) and even to a lesser extent (231Pa/235U). In the case of 226 Ra-excesses, three localities where documented eruptions have occurred (Section 3.14.3.2) provide key constraints, and we will discuss these in turn. A sample from the New Mounds site on the JDF Ridge has (226Ra/230Th) , 2.5 (Lundstrom et al., 1995), which is typical of many MORB (Figure 5(e)). Lundstrom et al. (1995) pointed out that when combined with other young samples from the area, there was in inverse

correlation of (226Ra/230Th) and (230Th/238U). However, the age constraints on the samples with lower (226Ra/230Th) in this data set are not good enough for this to be a firm conclusion. This general inverse correlation (rs ¼ 20.68, p ¼ 1) is also apparent in the larger database of “young” but not precisely dated samples we have compiled (Figure 11(a)), as has been pointed out by others (Kelemen et al., 1997; Volpe and Goldstein, 1993). The significance of the inverse correlation is significantly strengthened by recent data from the tube worm BBQ site ,98 500 N EPR (Sims et al., 2002) and the nearby Siquieros transform (Lundstrom et al., 1999). Although the samples from the Siqueiros transform are not well dated, these primitive depleted samples have extreme (226Ra/230Th) as high as 4.2, which are consequently minimum values. The samples from the 1991 –1992 New BBQ eruption can be usefully contrasted with the Siqueiros samples, forming a robust inverse correlation of (226Ra/230Th) against (230Th/238U). Significantly, a negative correlation is evident even in the more restricted range of (230Th/238U) of the definitive historic samples (Sims et al., 2002). These high precision analyses, from samples with very similar long-lived radiogenic nuclides, are thus a particularly well constrained data set and place critical constraints on the behavior of U-series nuclides during melting beneath ridges.

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Figure 11 (230Th/238U) versus: (a) (226Ra/230Th) for all fast and intermediate spreading ridge samples, dark gray symbols, and definitively historic eruptions, light gray symbols, (b) (231Pa/235U) for all samples with (226Ra/230Th) excesses (for which the effects of post-eruptive decay are a minimum), dark gray symbols, and definitively historic samples, light gray symbols. Different symbols also used to indicate different geographical regions. Despite a general lack of correlation, in some locations (230Th/238U) and (231Pa/235U) show some covariation.

There is no clear co-variation (rs ¼ 0.16, p ¼ 0.67) of (230Th/238U) with (231Pa/235U) for our compilation as a whole (Figure 11(b)). In keeping with earlier discussion, however, there is a sense of some co-variation at a local

scale. Sample from 98 300 –98 500 N EPR define a weak positive correlation (rs ¼ 0.48, p ¼ 0.97) whilst samples from JDF Ridge show a better but less significant correlation (rs ¼ 0.7, p ¼ 0.81).

U-series Melting Models Additional intriguing correlations are evident within the comprehensive, high-density geochemical data set of the ,98 500 N lavas (Sims et al., 2002). For example, these samples show correlations of Mg#, molar Mg/(Mg þ Fe), with 226 Ra – 230Th and 230Th– 238U disequilibria. Sims et al. (2002) strongly argue these trends results from variable mixtures of melts from different depths in the melting regime. Unfortunately, there are no other comparably detailed studies with which to compare these results, but this approach promises to be highly rewarding.

3.14.3.6 Summary of Observations The previous sections have described a number of key observations that need to be satisfied by successful models of melting and melt migration beneath ridges: . .

.

.

.

.

.

U-series disequilibria in young MORB are near ubiquitous. Most measured disequilibria are analytically significant and unlikely to result from shallow level contamination. Modest 230Th-excesses (,10 –20%) and large (,100%) 226Ra and 231Pa excesses are common in MORB unaffected by post-eruptive decay. On a ridge segment scale (230Th/238U) disequilibrium is frequently anticorrelated with (238U/232Th). This results in “local” linear arrays on equiline diagrams that represent a significant fraction of the global variability in (230Th/238U). (238U/232Th) can show significant variation on a local scale unrelated to 87Sr/86Sr and should not be used as a robust proxy of long-lived mantle heterogeneity. Historic samples show a negative correlation of (226Ra/230Th) with (230Th/238U), a trend further defined by less-well-dated samples. Apart from the 226Ra – 230Th excess anticorrelation, global variability of disequilibria with themselves and physical parameters are not well defined, perhaps masked to some extent by large local variability. – Local equiline arrays have variable slopes. These gradients do not show a clear variation with spreading rate as has been proposed (Lundstrom et al., 1998b), although 230Thexcesses show a weak, negative correlation with spreading rate. 231Pa-excesses show no spreading rate dependence. – (230Th/238U) shows a weak negative correlation with axial depth. – (231Pa/235U) shows a weak correlation with (230Th/238U) in two individual sample suites, but a negligible correlation for the global data set.

489

3.14.4 U-SERIES MELTING MODELS Geochemical and geophysical models are required that both explain the observations (Section 3.14.3.6) and provide better insight into the properties and processes occurring in partially molten regions beneath mid-ocean ridges. Because of the richness of the observations, this has proved to be reasonably difficult and this section reviews a range of models that have been proposed to explain U-series from the simplest closed-system models to fluid-mechanically consistent models of channelized melt transport. We compare model behavior to the geochemical systematics and discuss the inferences that are driven by the models. All of these models, however, require knowledge of the partitioning behavior of the U-series nuclides in mantle minerals (and some require information about their diffusivities) so we begin by discussing the constraints on these parameters from experimental petrology.

3.14.4.1 Preliminaries: Partition Coefficients and Diffusivity The initial reports of (230Th/238U) disequilibrium in MORB (Condomines et al., 1981) preceded reliable partitioning data for uranium and thorium. However, a plethora of ion-microprobe determined trace element partition coefficients has substantially changed this situation in the 1990s (Blundy and Wood, 2003a) and partition coefficients of uranium and thorium are now some of the best determined (Blundy and Wood, 2003b). A key feature is that all the nuclides of interest are highly incompatible in common mantle mineral phases (Table 2). Clinopyroxene and garnet (present in normal peridotitic mantle at depths greater than ,80 km) are the principal host minerals for uranium and thorium in the solid phase, although even in these phases partition coefficients do not exceed 0.05. An important consideration is the sense of uranium and thorium fractionation imparted by the presence of different minerals. This can be conveniently expressed in terms of DU/DTh. Minerals with DU/DTh . 1 retain uranium over thorium and contribute to 230 Th-excesses in a coexisting melt, whereas those with DU/DTh , 1 help to create 230Th-deficits. The effects of different minerals are then weighted by their absolute partition coefficients and modal abundance to control the bulk partition coefficient of the mantle, and thus determine the sense of fractionation of thorium from uranium in a coexisting melt. In spinel lherzolite, which likely comprises the main mantle lithology of the MORB melting regime, clinopyroxene largely controls

1,302

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Salters et al. (2002) 2.1 2.16 £ 1023 4.49 £ 1023 5.00 £ 1025 0.0007 3.80 £ 1025 0.0113

0.0057 0.028

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Salters et al. (2002) 1.3 8.92 £ 1024 1.18 £ 1023 5.00 £ 1025 0.0007 3.80 £ 1025 0.0015

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References Bulk Olivine Orthopyroxene Garnet Clinopyroxene Temperature (8C) Pressure, GPa

Table 2 Illustrative U and Th partition coefficients for common mantle minerals and bulk assemblages.

McDade et al. (2003)

Melt Migration in Oceanic Crustal Production Wood et al. (1999)

490

the fractionation of uranium and thorium. At “low pressures” (,1 GPa) thorium partitions preferentially into clinopyroxene relative to uranium. Early partition coefficient determinations of uranium and thorium in clinopyroxene documented such DU/DTh , 1 (Benjamin et al., 1980; LaTourrette and Burnett, 1992), and this was initially puzzling in the context of common 230 Th-excesses in MORB. Lundstrom et al. (1994) illustrated that large variations in the absolute values of partition coefficients of DTh, could be related to compositional variations in the clinopyroxene. This observation rationalized the large variations in DTh of the previous partitioning experiments as a result of different clinopyroxene compositions (e.g., LaTourrette and Burnett, 1992; Beattie, 1993a), but not the sense of fractionation. A further important effect on clinopyroxene partitioning is pressure. Crystal chemical changes with increasing pressure (notably decreasing CaO and increasing Al2O3) result in a decrease in the M2 site size, which hosts thorium and uranium. Thus the clinopyroxene increasingly prefers the smaller U(IV) relative to thorium (Wood et al., 1999). Thus, the sense of U –Th fractionation between clinopyroxene and melt changes with depth, such that between 1– 1.5 GPa, DU becomes greater than DTh (Landwehr et al., 2001; Wood et al., 1999). Garnet strongly incorporates uranium into its structure relative to thorium (LaTourrette et al., 1993; Beattie, 1993b; Salters and Longhi, 1999). The magnitude of DU and the ratio DU/DTh for garnet are larger than for clinopyroxene (Table 2) and so garnet dominates U – Th fractionation when present. Both the changing sense of U – Th fractionation in clinopyroxene and the eventual appearance of garnet result in “deep” melts having higher Th/U than the solid residue. Thus melting at depth is required to generate the correct sense of U – Th fractionation to account for 230Th-excesses in MORB. A more minor consideration is the role of orthopyroxene. At 1.5 GPa orthopyroxene has DU an order of magnitude less than clinopyroxene, and the small lattice site strongly favors uranium over thorium (Blundy and Wood, 2003a; Wood et al., 1999). The DU/DTh for orthopyroxene is consequently high (,2.5, Table 2) and may influence the bulk partition coefficient despite its low absolute partition coefficients, especially in highly depleted peridotites. The general picture outlined above is that the bulk DU =DTh , 1 for shallow mantle melting (,1 GPa) and so should result in 230Th-deficits (if the melt remains in equilibrium with the solid), whereas in the deeper mantle the converse is true. This overall picture seems secure. Some controversy persists, however, over the details of

U-series Melting Models the depth at which this change in sense of U –Th fractionation occurs and its magnitude. This is largely attributable to differences in the clinopyroxene partitioning data determined from experiments in high-pressure laboratories in Bristol (Landwehr et al., 2001; Wood et al., 1999) and Lamont (Salters and Longhi, 1999; Salters et al., 2002). Partition coefficients from the two different groups are reported in Table 2, using the two identical scenarios to facilitate direct comparison. Despite the ongoing debate, in a recent review of U-series partitioning, Blundy and Wood (2003b) point out some convergence between the parameterization of DTh and DU by Salters et al. (2002) and recent Bristol clinopyroxene partitioning data (McDade et al., 2003, accepted). There remain some differences for the “snap-shot” assemblages presented in Table 2, notably the lower overall clinopyroxene partition coefficient in the spinel stability field for the Salters et al. (2002) model. Salters et al. (2002) suggest that the aluminum contents of the clinopyroxenes of the Bristol group are unrealistically high, and that this may account for the high absolute values of their DTh and DU. On the other hand the 1.5 GPa experiments of Salters et al. (2002), which form an important part of their parametrization, were conducted at temperatures (e.g., ,1,400 8C at 1.5 GPa) well in excess of most estimates of the mantle solidus (Hirschmann, 2000). Such elevated temperatures could account for anomalously low partition coefficients. It would be useful to resolve these differences, as the models presented below (Section 3.14.4.3) are quite sensitive to the values of partition coefficients used (e.g., see Jull et al., 2002). It should also be stressed that the values cited in Table 2 are only illustrative values for specific fertile mantle compositions at the onset of melting. As melting continues, the compositions and modes of minerals change as the peridotite becomes less fertile. Thus, mapping out the variation in bulk distribution coefficients is a complex task (Salters et al., 2002), and the values cited should not be treated as absolute values for the MORB melting regime. Fractionation of the other nuclide pairs of interest (226Ra/230Th) and (231Pa/235U) rely in part on the partition coefficients of thorium and uranium discussed above. The partition coefficients of radium and protactinium have not been directly determined but can be estimated from applying a lattice strain theory to extrapolate data from homovalent cations (Blundy and Wood, 1994, 2003b). This approach suggests that radium and protactinium are indeed highly incompatible in the likely MORB source mineral assemblage, as is inferred empirically from the large degrees of disequilibrium. In both cases, the daughter nuclides are likely more than an order of magnitude more incompatible than their short-lived parent

491

(Cooper et al., 2001; Blundy and Wood, 2003a). In this case, the actual value of the distribution coefficient of the daughter nuclide is not critical and fractionation is then dominantly controlled by the value of the less incompatible parent, which are fortunately much better constrained. For once, Nature smiles on us.

3.14.4.1.1

Diffusion coefficients for U-series nuclides

Use of partition coefficients alone to model element fractionation assumes a melt process controlled by equilibrium partitioning. If melting is rapid relative to the diffusion of elements through mantle minerals, equilibrium may not be maintained (e.g., Qin, 1992; Van Orman et al., 2002a) and this has been invoked to be a potential cause of U-series disequilibrium in MORB (e.g., Rubin and Macdougall, 1988; Volpe and Goldstein, 1993). It is, thus, also important to consider U-series element diffusivities. Seitz (1973) first presented data on uranium and thorium diffusion in clinopyroxene. Recently, the diffusivities of thorium and uranium in clinopyroxene have been redetermined and are so low (,3 £ 10221 m2 s 21 at 1,200 8C) to suggest equilibrium is unlikely during MORB melting (Van Orman et al., 1998). However, differences in diffusivity between thorium and uranium are small and in a sense that should produce minor 238U rather than 230 Th excesses during melting (Van Orman et al., 1998). Application of an elastic diffusion model (Van Orman et al., 2001) for divalent cation diffusivities suggests that radium should diffuse several orders of magnitude more rapidly than thorium, and so this can potentially account for 226Ra-excesses in a nonequilibrium model. Predicting the likely opposing effects of smaller ionic radius but larger ionic charge of Pa5þ relative to U4þ is difficult, but it seems improbable that procactinium diffusivities are sufficiently large to explain 100% 231Paexcesses by disequilibrium melting. Unfortunately, there are no directly measured data for uranium and thorium diffusivity in garnet, and so full characterization of the state of equilibrium during MORB melting remains to be determined.

3.14.4.2 Implications for Closed-system Models The very small bulk partition coefficients for the U-series nuclides in the mantle ðD ,, 0:01Þ puts a major constraint on the viability of simple closed-system models (Section 3.14.2.2). In these models, element fractionation is only efficient when the degree of melting F is comparable to the bulk partition coefficient. The degree of melting is

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Melt Migration in Oceanic Crustal Production

not well known in many melting environments. In the MORB case, however, the thickness of the oceanic crust coupled with an extensive set of melting experiments constrains melting beneath normal ridges to be ,10% (e.g., Klein and Langmuir, 1987; McKenzie and Bickle, 1988; see Chapter 3.13). Such large amounts of melting would not be expected to produce any U-series disequilibria from closed-system models. Because Osmond (2003) assumed that MORB melting was governed by such a model, he was forced to conclude that there were highly refractory accessory minerals in the mantle that retained uranium and thorium with Ds comparable to the degree of melting in MORB. This inference is entirely model dependent, however, and is easily relaxed if one considers open-system models (which are still based on fundamental statements of conservation of mass for radioactive tracers).

3.14.4.3 Open-system “Ingrowth Models” The closed-system models are not sufficient to explain U-series excesses in MORB; however, they form a useful precursor to open-system models that allow the melt to move relative to the solid. Melt separation and extraction must occur at some scale for volcanism to be observed. Once melt can move, however, additional processes come into play that can produce significant radioactive disequilibrium, even at large degrees of melting. These models are often described as “ingrowth models” because they produce excess daughter products over that in the unmelted source by ingrowth from the parent. Spiegelman and Elliott (1993) provide a discussion of many of these models and stress that the key to understanding their behavior is that radioactive disequilibrium will arise from any process that produces different residence times for parent and daughter nuclides in the melting zone. As an extreme example, consider an open melting system where the parent is completely compatible in the solid and is never extracted while the daughter partitions completely into the melt. The residence time of the parent in the system would be the time it takes the solid to cross the melting regime. The residence time of the daughter, however, would be the melt extraction time (i.e., once it is produce it leaves the system in the melt). During the extra time the solid spends in the system; however, it is continually decaying and producing extra daughter product, which is observed in the melt. If the melt is continually allowed to separate from the system, it will contain only daughter product and have an infinite excess. This process can be kept up indefinitely as long as the radioactive production can keep pace with extraction. In this case, the excesses are

controlled by the rate of melting and extraction and less by the total amount of melting F. At the other extreme, if the unmelted solid is in secular equilibrium at the initiation of melting and all elements spend the same account of time in the system, then they will remain in secular equilibrium because they continue to all decay at the same rate. The important (and confusing) feature of radioactive decay chains is that these systems do not have to conserve mass for any individual nuclide because it can always be produced by the parent or decay depending on the relative concentration of parent and daughter nuclides. For example, there is always less of the principal parent (e.g., 238U) leaving the system than entering it, simply because of the radioactive decay of the parent over the timescale of melt extraction. For a long-lived nuclide like 238U the resulting differences in concentration are negligible. However, for short-lived nuclides like 226 Ra, the actual 226Ra atoms that leave the system may not even be the same ones that entered. Any initial 226Ra in the melt can decay and be replaced by later production or “ingrowth” from longer-lived 230Th. For further discussion, see Spiegelman and Elliott (1993) or Elliott (1997). The following sections consider a host of ingrowth models and discuss how each one produces radioactive disequilibria. We first discuss the different mechanisms inherent in the models and then compare their behavior to each other and to the observations.

3.14.4.3.1

“Dynamic melting models”

All the models in this section are “solid-centric” in the sense that they track the behavior of a small parcel of solid undergoing melting. In the equilibrium dynamic melting models, this solid is assumed to be in equilibrium with a small fraction of melt in contact with the solid. The diffusion limited models simply relax the assumption of complete chemical equilibrium between solid and retained melt. The local system of solid plus retained melts acts much the same as the closed-system models described in Section 3.14.2.2.2; however, the excesses are controlled by the retained melt fraction f0 (usually assumed to be the porosity) relative to the partition coefficients ðf0 =DÞ: This differs from the closedsystem models where the porosity is assumed to be equal to the degree of melting F and the excesses are controlled by F=D: The important feature of all open-system models is that the porosity must be significantly smaller than the degree of melting if the melt is to separate from the solid (see below).

U-series Melting Models In the dynamic melting models, all melt fractions produced in excess of f0 are assumed to be mixed and extracted, although the dynamics of these processes are not included in the models. For nonradiogenic elements, these models are identical to commonly used dynamic melting models for trace elements (e.g., Langmuir et al., 1977; Johnson and Dick, 1992; Sobolev and Shimizu, 1993; Slater et al., 2001; see Chapter 3.13). The principal differences between all of the dynamic ingrowth models are the degree of equilibration maintained between solid and retained melt and how the extracted melts are integrated. Equilibrium models. Equilibrium dynamicmelting U-series models were first introduced by McKenzie (1985) and Williams and Gill (1989). For stable elements, these models are all identical to “near-fractional” melting models (e.g., Johnson and Dick, 1992). For U-series elements however, they can produce significant radioactive disequilibrium by ingrowth if the melting rate is slow with respect to the decay rate of the short-lived daughter. As long as there is sufficient time for new daughter to be produced and concentrated in the equilibrium porosity before it is extracted, then reasonable excesses can be produced throughout the melting process. In the limit of extremely slow melting, the maximum excess in the retained melt (and, therefore, the accumulated melt if there is no time for decay) is ! 230 Th ð1 2 f0 ÞDU þ f0 ¼ ð8Þ 238 U f ð1 2 f0 ÞDTh þ f0 Therefore, these models put constraints on the retained porosity f0 in the model relative to the partition coefficients. If melting is much faster than ingrowth, then these models produce accumulated near-fractional melts that reflect the source concentrations (and therefore will be close to secular equilibrium). Williams and Gill (1989) actually present several equilibrium models. Their “dynamic melting” model is the same as McKenzie’s but they also introduce an “accumulated continuous melting” (ACM) model. The only difference between these two models is how the incremental or “continuous” melts are pooled and extracted. The accumulated continuous melting model is somewhat easier to understand. It follows a piece of solid through the melting process and simply accumulates the extracted melt in a chemically isolated reservoir. Any excesses in this reservoir are unsupported and decay toward secular equilibrium. Thus, there is a trade-off between melting rate and excesses in this model. If the melting rate is too slow, the time required to melt to degree F is long compared to the half-life of the daughter nuclide and the accumulated melt will have

493

negligible excesses. Too rapid melting also leads to secular equilibrium because both parent and daughter are stripped from the solid by chemical fractionation before any ingrowth can occur. It is reasonably difficult to produce significant excesses (particularly radium) using this model. In contrast, the dynamic melting model assumes a continuous “melting column” appropriate for adiabatic decompression melting. This column is composed of a series of solids all in different stages of melting from F ¼ 0 at the bottom to F ¼ Fmax at the top, with a melting rate proportional to the solid upwelling rate. In this model, the incremental melts from all depths in the melting column are mixed and assumed to be extracted instantaneously. This model produces significant radioactive disequilibrium because none of the excesses in the accumulated melt have time to decay. As emphasized in Spiegelman and Elliott (1993), most of the excesses in this model are produced near the bottom of the melting column because these melts are most enriched and because phases capable of fractionating thorium from uranium are more prevalent at higher pressures. This model can produce U-series excesses comparable in magnitude to those that are observed in MORB if the retained porosity near the bottom of the melting regime is small compared to the partition coefficients. Because this model assumes instantaneous extraction, all excesses produced by the calculations should be treated as upper bounds. The dynamic melting model was extended to two-dimensional systems by Richardson and McKenzie (1994). In their model, each solid flow line is treated as a single dynamic melting column and the melts from the entire region are mixed and extracted instantaneously. This model tends to produce larger excesses than the single column model because it is dominated by smaller degree, slower melting regions from off-axis. How these melts actually migrate to the ridge axis and whether they do so without losing their radioactive disequilibrium is unclear because melt transport is not included in these models. An interesting feature of their simplified ridge model, however, is that the excesses produced are nearly independent of spreading rate. The argument is best developed in Iwamori (1994), who assumes a passive solid flow field beneath the ridge driven by a thickening wedge-shaped lithosphere that is thicker for slow-spreading ridges than fastspreading ridges (e.g., see Spiegelman and McKenzie, 1987). Thicker lithosphere drives a faster upwelling which offsets the reduction in spreading rate such that the actual upwelling rate (and, therefore, melting rate) is reasonably constant at all spreading rates. They offer this as an explanation for the relatively weak (or negligible)

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Melt Migration in Oceanic Crustal Production 230

correlation between Th-excesses and spreading rate (Figure 9). More recently, Zou and Zindler (2000) provide a tractable analytic solution to both McKenzie’s and Williams and Gill’s equations for the limited problem of constant partition coefficients and porosities. By making all adjustable parameters constant, these models exploit the analytic solutions afforded to linear systems of ordinary differential equations like Equations (1) – (3) (a similar approximation for the transport models described in Section 3.14.4.3.2 can be found in the appendix of Spiegelman and Elliott (1993)). However, because the solutions of these models are sufficiently sensitive to parameters such as D=f0 it can be challenging to choose specific constants that are actually appropriate to a problem where the properties of the system are expected to change in space (see Section 3.14.4.1). Disequilibrium/Diffusion limited models. The principal parameters controlling solutions to the equilibrium “dynamic melting models” are the melting rate with respect to the half-life of the daughter product ðG=½rs t1=2 Þ and the retained porosity with respect to the partition coefficient ðD=f0 Þ: Given the very small bulk partition coefficients expected for the U-series nuclides ðD p 0:01Þ; these models place significant constraints on the retained porosity f0 : As pointed out by Spiegelman and Elliott (1993), f0 is not necessarily the physical melt fraction present, rather it is the fraction of melt that remains in equilibrium with the solid. Nevertheless, these models require that this number is very small, i.e., comparable to the partition coefficients of the U-series nuclides. Using a similar model for stable elements to estimate f0 from the depletion of abyssal peridotites give f0 , 1% (Johnson et al., 1990; Johnson and Dick, 1992) (which would produce no U-series excesses). Similarly, Slater et al. (2001) require f0 , 3 – 4% to model the variability of melt-inclusions in Iceland. The inconsistencies between inferred porosities given different chemical systems suggests either different models are required or perhaps that strict equilibrium partitioning needs to be relaxed for the U-series. One possibility that would lead to larger inferred porosities for the U-series was introduced by Qin (1992, 1993), who proposed that the retained melt was only in complete equilibrium with the surface of minerals and that solid-state diffusion limited the re-equilibration of the retained melt with the solid. In other respects, this model is identical to the ACM model of Williams and Gill (1989). Qin introduced a specific microscopic melting/diffusion model for spherical grains and coupled it to the larger-scale dynamic melting models. The net affect of this

addition was to drive the effective partition coefficient of all elements toward 1 so melting produced less fractionation between elements. Again there is a trade-off between melting rate and diffusion rates. If the melting rate is rapid with respect to diffusion, the problem reduces to disequilibrium melting of the particular phase and the melts will simply pick up the activity ratio of the solid phase (e.g., as in Section 3.14.2.2.2). Slow melting or rapid diffusion returns the equilibrium dynamic melting models. However, beyond changing the effective partition coefficient and potentially allowing larger porosities, the disequilibrium models have the same mixing and transport assumptions as the equilibrium models and, therefore, similar inferences and caveats.

3.14.4.3.2

Single porosity transport models

The physics of melt extraction and mixing are not included explicitly in the dynamic melting models. However, melts must move faster than the upwelling solid, otherwise melt separation would not occur, and the system would reduce to pure batch melting to degree of melting Fmax (which produces no excesses at large F). By explicitly including the production and transport of melt, transport models of U-series excesses add an additional velocity field that can affect the overall residence times of parent and daughter nuclides as well as affect the mixing of melts. These models also produce excesses for large degrees of melting and place constraints on the rates of melting and extraction processes. However, the inferences drawn from these models can be quite different from the dynamic melting models. The simplest transport model was introduced by Spiegelman and Elliott (1993) and consists of a one-dimensional, steady-state melting column where it is assumed that the melt and solid remain in chemical equilibrium at all times. These models are readily extended to spatially variable melting functions FðPÞ and partition coefficients DðPÞ (Lundstrom et al., 1995) and a more user friendly web-based version of these models is available at http://www.ldeo.columbia.edu/ ~mspieg/UserCalc/. The web-site and U-series calculator is described in detail by Spiegelman (2000). Figure 12 shows some output from UserCalc for a one-dimensional upwelling column and the bulk partition coefficients given in Table 2. The melt velocity in these models at any height z in the column is of order wðzÞ < W0

FðzÞ fðzÞ

ð9Þ

where W0 is the solid upwelling velocity, FðzÞ is the degree of melting at height z, and fðzÞ is

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0.2

10–5 10–4 10–3 10–2 0 Bulk D

1

2

3

4

5

Activity ratios

Figure 12 Example output of UserCalc (Spiegelman, 2000) using bulk partition coefficients from Table 2. Physical parameters for these calculations are a melting column of height 90 km, an upwelling rate W0 ¼ 7 cm yr21, a maximum degree of melting Fmax ¼ 0.1 and a maximum porosity fmax ¼ 0:002: These parameters give a maximum melt velocity that is 50 times faster than the upwelling velocity. This permeability is consistent with estimates from Von Bargen and Waff (1986). Both calculations assume a two-level partition coefficient scheme where the “Garnet stability field” extends from 3 – 2 GPa and spinel is stable for pressures shallower than 2 GPa. (a) Calculations using Bulk Ds from Table 2 after Salters et al. (2002). DRa ¼ DPa ¼ 1025 : (b) Bulk Ds from Table 2 after McDade et al. (2003) (garnet) and Wood et al. (1999) (spinel). The first panel is the same in both cases and just shows degree of melting F(z) and porosity fðzÞ as a function of pressure. The second panel shows the bulk-partition coefficient structure with pressure and the final panel shows activity ratios for (230Th/238U), (226Ra/230Th), and (231Pa/235U). The activity ratios in (a) are comparable to those observed for 9 –108 N EPR although the 226Ra-excess is too small. The activity ratios in (b) are low for (230Th/238U) but quite large for Ra and Pa. This results from the large DU and DTh in the upper region. Note the large increase of Pa due to ingrowth in the upper layer of (b). These plots illustrate the effect of partition coefficients on modeling U-series but assume a rather simple D structure. More realistic models should include more variable D with depth that accounts for pressure and changing mineral modes.

Melt Migration in Oceanic Crustal Production

where D0 ¼

rs ð1 2 fÞDi D < i rf f f

ð11Þ

is the effective partition coefficient and depends (again), principally on D=f (rs and rf are the densities of solid and melt, respectively). Elements with partition coefficients much smaller than the porosity travel near the melt velocity, while those with larger partition coefficients travel more slowly. As discussed in Spiegelman and Elliott (1993), these models have two regimes, one where the excesses depend on the differences in effective velocities, the other which depends on the ratio of effective velocities. While the details of these models are somewhat complicated (see Figure 13), their essence is straightforward. Any region where the parent nuclide moves more slowly than the daughter will cause ingrowth and excess activities of the daughter.

(230Th/238U) Porosity coupled to upwelling rate

1. 2 1.1 5 1.1

10

1.3

the physical porosity at that same point. For example, if the melt velocity is 50 times faster than the upwelling solid velocity at the top of the column ðFmax ¼ 0:1Þ; then it requires that the porosity in this region to be 50 times smaller than the degree of melting or fmax ¼ 0:002: When all the melt extraction is through this single porosity, conservation of mass simply requires that the faster the melt moves, the less melt can be retained. While a porosity of 0.2% seems small, it is consistent with estimates of permeability by Von Bargen and Waff (1986) that give the appropriate velocities at this porosity for a grain size of ,3 mm (see Spiegelman and Kelemen, 2003, for details). If the porosities are also small with respect to the partition coefficients of the U-series nuclides, these models can also produce significant U-series disequilibrium. Figure 12 shows calculated porosity structure and activity ratios for a typical problem. For stable elements or nuclides with half-lives much longer than the melt extraction times (such as 238U and 232Th), the melts that leave the top of this column have concentrations equivalent to that for a batch melt with degree of melting Fmax independent of porosity (see Ribe, 1985; Spiegelman and Elliott, 1993). Therefore, these models cannot fractionate highly incompatible elements or produce dispersion on an equiline plot. They can, however, produce large U-series excesses if the parent spends longer in the system than the daughter. As discussed at length in Spiegelman and Elliott (1993), the residence times are controlled by the effective velocity of parent and daughter elements which is approximately w þ WD 0 wieff ¼ ð10Þ 1 þ D0

Upwelling velocity

496

1

0.1 0.001

0.01 Porosity

Figure 13 Possible variability of (230Th/238U) as a function of spreading rate. This figure shows contours of (230Th/238U) as a function of upwelling rate and maximum porosity (i.e., melt velocity) from UserCalc. These contours are the activity of the top of the melting column for the model shown in Figure 12(a) (see Spiegelman and Elliott (1993), for a detailed description of these plots). If the porosity were independent of upwelling velocity (dashed arrow), this specific calculation model would predict increasing then decreasing (230Th/238U) with increasing spreading rate. However, for a fixed permeability structure, faster upwelling implies larger porosities. The solid arrow shows a possible set of consistent upwelling rates and porosities for a series of runs for 1 # W0 # 10 cm yr21. Along this trajectory, variations in (230Th/238U) are small.

One of the interesting features of these models is that observable excesses for different parent/daughter pairs can be produced in different parts of the melting region. For example 230Th has a sufficiently long half-life such that excesses produced at the bottom of the column in the garnet stability field can be preserved at the surface. Unless melt transport is extremely fast, however, excess 226Ra produced at the bottom will decay away before eruption. Nevertheless, significant 226 Ra can be produced near the top of the melting column if the thorium and radium velocities are sufficiently different (e.g., Figure 12(b)). Thus thorium excesses and radium excesses can be decoupled. If 226Ra is observable at the surface, it is really providing information about the 230Th that produced it within the last ,8 kyr. In the dynamic melting models, that 230Th is at the bottom of the melting zone and requires rapid extraction to preserve it. In the equilibrium transport models, the parent 230Th is near the surface. However, for these models to produce large 226Ra-excesses requires that the porosity near the surface is comparable to DTh :

U-series Melting Models Disequilibrium single porosity models. In addition to complete equilibrium transport, several other variations to the basic model have been proposed. The first relaxes the assumption that moving melt remains in chemical equilibrium with the solid at all times (Spiegelman and Elliott, 1993), although instantaneous melts are assumed to be in chemical equilibrium with the mantle that produced them. For stable elements, this disequilibrium transport produces a residue that reflects perfect fractional melting and the melts have compositions identical to accumulated fractional melts. These models are similar to the dynamic melting models in the limit that f0 ! 0 but include the time it takes the melt to leave the system. These models produce most of the disequilibrium at the bottom of the melting column and produce negligible U-series excesses due to decay during transport. It is also straightforward to extend the equations to allow for only partial equilibration during transport. Iwamori (1993a) presents a one-dimensional steady-state single-porosity model for stable elements that includes diffusive re-equilibration between melt and solid. He does not extend it to radioactive nuclides in this paper but includes this effect in his “two porosity model” (Iwamori, 1994) (see Section 3.14.4.3.4). The expected effects of chemical disequilibrium should be similar to those in the Qin (1992) dynamic melting model, namely he effective bulk partition coefficients of all elements will be driven towards unity. Linear kinetic approximations to melt/solid diffusion are also possible, but all of these models should simply produce results that are intermediate between full equilibrium and disequilibrium transport. In general, all of the transport and “dynamic” melting models produce a single melt composition at the top of the melting column for stable elements that has highly incompatible trace element ratios that are indistinguishable from the source. 3.14.4.3.3

Comparison with observations

At this point, it is worth comparing the behavior of both the dynamic melting models and the transport models to each other and to the key observations listed in Section 3.14.3.6 to see how well these models explain the observations and to suggest that more complicated models may still be needed. Both dynamic melting and transport models can produce the magnitude of observed excesses for the large degrees of melting expected at mid-ocean ridges. However, these excesses are produced in different parts of the melting regime and have different implications for the rate of melt transport. The dynamic melting models produce most of their excesses near the bottom of the

497

melting column and require rapid melt extraction (on timescales short with respect to the half-life of 226 Ra) to preserve them at the surface. The equilibrium transport models however, produce their excesses throughout the column and in particular can produce the 226Ra-excesses near the surface thus reducing some of the requirement for fast melt transport. Nevertheless, these models still require that the porosity near the top of the column is small with respect to DTh and this indirectly puts constraints on the melt velocity. To produce significant 226Ra-excesses by transport, requires that the porosity f is comparable to DTh. In the single-porosity model, however, the entire melt flux is required to pass through this porosity which implies a large melt velocity via Equation (9) (which is strictly valid only for onedimensional, steady-state melting columns with a single fðzÞ). For example, if the degree of melting Fmax ¼ 10% and DTh , 0:001 near the top of the melting column, then the melt velocity will be ,100 times larger than the upwelling velocity or 1 –10 m yr21 to produce significant 226 Ra-excesses. These rates are not excessive, however, and can easily be attained by porous flow using the parametrization of Von Bargen and Waff (1986). Spiegelman and Kelemen (2003) provide an example and details. The melt velocity estimated from transport models is a bit slower but still comparable to estimates from the dynamic melting models. For example, if we assume that 226Ra-excesses are produced at the bottom of a column 90 km deep and need to move to the surface in ,3 halflives, then w0 , 20 m yr21. It should be stressed that this is a constraint on the average melt velocity across the entire melting column rather than a constraint on the maximum velocity near the surface. Moreover, the constraint from Equation (9) assumes that there is only a single porosity near the surface. Two-porosity models (next section) relax this constraint somewhat. Figure 12 gives a sense of the magnitude of 226 Ra-excesses possible for a single column with Fmax =fmax , 50 ðfmax ¼ 0:002Þ but with different partition coefficients. In Figure 12(a), DTh ¼ 8:9 £ 1024 , fmax and produces small 226Raexcesses ((226Ra/230Th) , 1.6). Figure 12(b), however, has DTh ¼ 5:6 £ 1023 . fmax ; which produces a significantly lager excess ((226Ra/230Th) , 4.3). For simply producing the magnitude of observed excesses, both dynamic melting models and transport models are viable. The largest uncertainty in either type of model is the appropriate partition coefficients (and D as a function of pressure) as these control the inferences on the porosity through the parameter D=f: The more stringent tests on these models, however, come from the observed correlations

498

Melt Migration in Oceanic Crustal Production

between nuclides and excesses. In these tests, both sets of models have difficulty which has driven a need for more complex models. We begin by discussing the behavior of these models on an equiline diagram and compare it to the observation that individual ridges appear to produce correlated arrays on these diagrams, even for an isotopically similar source. In their usual forms, both dynamic melting models and transport models simply cannot fractionate highly incompatible elements for degrees of melting F . DU ; DTh (as is expected at ridges) if the melts that are “erupted” from the model are assumed to be the melts from the top of the column. For dynamic melting models, this is because usually all of the incremental melts are assumed to be efficiently mixed. Thus, for stable elements, the melt composition is similar to an accumulated fractional melt which has ratios of highly incompatible elements such as [U/Th] that are close to source. Similarly, the equilibrium transport models produce melts with compositions identical to batch melts (which are hardly distinguishable from accumulated fractional melts for elements with D p F). Thus both models produce final erupted melts with (238U/232Th) very close to the source ratios, which implies only vertical motion on an equiline diagram. That is, if the unmelted source lies on the equiline at a given (238U/232Th), then the final melt moves to a higher (230Th/232Th) but at roughly the same (238U/232Th). Simple ingrowth models, with complete mixing, cannot produce an array on an equiline diagram without invoking source variability which was done Lundstrom et al. (1998a,b, 2000) to try to explain the local arrays. Nevertheless, one of the key observations for at least 98 N EPR and the Reykjanes Ridge is that individual arrays can show significant variations in (238U/232Th) for uniform 87Sr/86Sr suggesting that this variation cannot be entirely attributed to source. It is possible for the basic ingrowth models to produce arrays, however, if one allows incomplete mixing over the full range of melt compositions. For example, for dynamic melting models, if the incremental melts are not assumed to mix, they form a very large array of melt compositions from very enriched small-degree melts to extremely depleted fractional melts. For the very small porosities required to produce the U-series excesses, however, the range in concentrations and the stable trace element patterns produced by these models are quite different from what is observed (e.g., see Slater et al., 2001). For transport models, if the full trajectory of melts from the bottom of the melting column to the top are plotted on an equiline, this will produce an array of melt compositions (e.g., see figure 4 in Spiegelman and Elliott, 1993). However, since these melts are automatically mixtures of all melts

produced up to that height, they very rapidly attain (238U/232Th) ratios comparable to source for any F q DU ; DTh. Moreover, the most depleted melts are the mixed melts at the top of the column. Thus, this model can never produce a melt more depleted than a batch melt with degree of melting Fmax. The instantaneous “continuous” melts from the dynamic melting model, however, can become extremely depleted. Another problem with the transport models is that they assume complete equilibrium between melt and solid throughout the melting regime. This implies that the solid residues at the top of the column (e.g., near the Moho), should be in chemical equilibrium with MORB. However, another key observation of MORBs is that they are out of equilibrium with abyssal peridotities near the Moho for both major and trace elements. A quick fix for the transport models is to assume that melts remain in chemical equilibrium up to some depth and then melt fractionally for the remaining distance (e.g., see Kelemen et al., 1997). However, different attempts to explain both U-series and stable elements in melts and residues in the single-porosity transport models has motivated much of the development of the two-porosity models in next section. The basic models also have difficulties accounting for the observed correlations (or anticorrelations) among radioactive excesses. In the case of the observed anticorrelation of (226Ra/230Th) versus (230Th/238U) at intermediate to fastspreading ridges (Figure 11), the dynamic melting models have considerable difficulty because all U-series excesses are produce in the same narrow region at the bottom of the melting regime. This does not necessarily rule the model out because variations in melting rate can cause changes in 230 Th without affecting 226Ra if the melting rate is slow with respect to radium decay, but fast with respect to thorium (e.g., Spiegelman and Elliott, 1993; Figure 7(a)). However, it is still difficult to produce an anticorrelation using these models. Moreover, for these variations to be observed in closely spaced samples would suggest a melting regime where the solid velocity varies dramatically over small length scales. So far, no solid flow model has been developed with this property. As for the transport models, because most 230 Th-excesses are produced at the bottom of the column while 226Ra is produced near the top, it is possible to decouple them. Nevertheless, for many systems this model still produces positive correlations. For example, all excesses will tend to increase with decreasing porosity. However, it is possible to produce large 226Ra-excesses with low (230Th/238U) if a long time is spent in the spinel field, while low (226Ra/ 230Th) and high ( 230Th/ 238U) implies more rapid extraction from depth. The problem is how to produce

U-series Melting Models the anticorrelation between the two in closely associated samples in a systematic way with these models (Figure 11). Similar arguments can be made for the apparent lack of correlation of (231Pa/235U) with (230Th/238U) if the large 230Th-excesses are coming from depth and the 231Pa-excesses are being ingrown from shallower parts of the melting column. Nevertheless, one might expect a positive correlation between these two systems from the models because parameters that increase 230 Th ingrowth by retaining 238U should also increase 231Pa ingrowth by retaining 235U. More data probably need to be collected to explore the possible positive correlation between (231Pa/235U) and ( 230Th/ 238U) hinted at in some local assemblages. Perhaps one of the more surprising observations from U-series studies is the weakly negative, to possibly nonexistent correlation between spreading rate and U-series excesses (Figure 9). In the simplest dynamic melting models one might expect a direct correlation because all excesses are produced in a narrow region and are directly influenced by melting rate which is assumed to scale with spreading rate. In these models, independent of source, slower melting produces larger excesses which is only hinted at in the date. For transport models, the problem is slightly more complicated because upwelling rate affects both the melting rate and porosity (melt velocity). Figure 13 shows contours of (230Th/238U) at the top of the model in Figure 12(a) as a function of porosity and upwelling rate as calculated by UserCalc (see Spiegelman and Elliott, 1993, for a detailed discussion of these contour plots). If porosity is independent of upwelling rate, then one might expect to see variations in (230Th/238U) with upwelling rate but they need not be monotonic. The dashed arrow in Figure 13 shows a trajectory through porosity-upwelling velocity space for constant porosity ðf0 ¼ 0:002Þ and 1 # W0 # 10 cm yr21 : For this particular run, this trajectory first shows an increase in (230Th/238U) that is a consequence of spending less time in the shallow spinel field followed by a decrease in (230Th/238U) from the melting rate exceeding the ingrowth rate. However, if the functional form of the permeability is independent of upwelling rate, then faster upwelling (faster melting) systems should produce larger porosities. A scaling argument for the maximum porosity as a function of upwelling rate is given in Spiegelman (2000), which for high-permeability systems is approximately

fmax / W01=n

ð12Þ

when n is the exponent in the permeability/ porosity relationship which is approximated in this model by kf ¼ k0 fn : Such a coupled set of velocities and porosities is shown in Figure 13 by

499

the solid arrow. For this calculation, at least, the contours of activity are subparallel to the arrow suggesting that the overall variation of 230 Th-excesses would show little change over an order of magnitude spreading rate variation which is consistent with the data. More specifically, this particular trajectory suggests that 230Th-excesses should be reasonably constant (or slightly increasing) for slow to intermediate upwelling rates and only show a noticeable decrease for faster upwelling. This qualitatively agrees with the variation of 230Th-excesses shown in Figure 9 but it could be entirely coincidental. The final test of the models is against axial depth (Figure 10). This issue is discussed in detail by Bourdon et al. (1996a) using a smaller data set from the Azores platform (FAZAR cruise). Again, in their simplest form, the dynamic melting models would suggest that 230Th-excesses should be nearly independent of axial depth. If we assume that axial depth reflects the total amount of melt production and, therefore, the depth of the melting column, then the simplest dynamic melting model produces all excesses in a narrow region near the bottom. If the partition coefficients do not change significantly within such a small depth range, then all dynamic models with instantaneous transport should produce roughly the same excesses for all columns that start in the garnet field. The only systematic change would be for melts that begin melting in the spinel field. This argument needs to be modified somewhat by the recent results on the pressure dependence of partition coefficients for the U-series (e.g., Wood et al., 1999). If in general, DU =DTh is an increasing function of pressure, then it is possible that dynamic melting models would show a weak positive correlation with melt production (or negative correlation with axial depth). If the actual transit time is factored in, however, then it is possible that these models would produce smaller 230Th-excesses with larger crustal thickness because the deeper produced melts have further to travel and will decay during transport. Equilibrium transport models, however, should show a negative correlation of 230Thexcess with axial depth (assuming that axial depth is a proxy for depth of melting) because deeper melting columns imply a longer time spent in the garnet stability field and higher degrees of melting tend to produce larger porosities and, therefore, faster transport through the spinel field. Bourdon and Langmuir (1996) show this calculation using the equilibrium transport model of Spiegelman and Elliott (1993) for a series of melting columns with constant upwelling rate but increased depth of melting. The calculation can become more complex, however, if upwelling rate is somehow coupled to depth of melting. In summary, both the dynamic melting model and the transport models can usefully explain

500

Melt Migration in Oceanic Crustal Production

some subset of the observations but remain problematic for explaining all of them. To move forward, many authors have developed hybrid models that combine aspects of both dynamic and transport models by invoking some form of channelized melt flow. These “two-porosity models” vary in degree of complexity but have promise for reconciling many of the observations. We now discuss them in detail.

3.14.4.3.4

Two-porosity models

A fundamental observation of melt transport at mid-ocean ridges is that in both major and trace elements, MORBs are out of chemical equilibrium with the shallow upper mantle (e.g., see Kelemen et al., 1997, for review). One mechanism that has been suggested is that melt migration is localized into some form of channel network fed by porous flow at the grain scale. This basic idea has been incorporated into several different “two-porosity models” to try to reconcile U-series observations with those of stable trace elements. The first model combining porous and “channel flow” was developed by Iwamori (1993a,b) for stable elements and then extended to U-series elements (Iwamori, 1994). This model combines aspects of the full transport models with the instantaneous melt extraction of the dynamic melting models. In some sense, this model is a more complex version of the dynamic melting models where the porous flow regime produces excesses that are extracted into chemically isolated channels, mixed completely and extracted instantaneously. That is, like the dynamic melting models, the actual time spent in the channels is not included in the calculations of excesses. However, the porous ingrowth regime is a full transport model with diffusion controlled re-equilibration, rather than a simple constant-porosity regime that moves with the solid. The relative amount of porous and channel flow is controlled by a “suction parameter,” S, such that S ¼ 0 is essentially the equilibrium transport model of Spiegelman and Elliott (1993) and S ¼ 1 is the purely fractional dynamic melting model of McKenzie (1985) in the limit f0 ! 0: This model is a continuous hybrid of all the previous models. Its principal weakness is that it still assumes instantaneous extraction and mixing in the channels and thus the channel melts provide an upper bound on the excesses. Lundstrom (2000) remedies this problem by assuming the channels are chemically inert dunites undergoing porous flow at a different porosity as suggested by Kelemen et al. (1995a) (see also Kelemen et al., 1995b, 1997; Kelemen and Dick, 1995; Aharonov et al., 1995). The dunites are assumed to occupy a fraction x of the

whole system and the flux into the channels is controlled by the same suction parameter as in Iwamori (1994). This model was principally developed to try to reconcile the important observation that abyssal peridotites are extremely depleted, implying near-fractional melting (Johnson et al., 1990; Johnson and Dick, 1992; Hellebrand et al., 2002), while U-series excesses, particularly short-lived excesses of 226Ra and 231Pa, are more easily made with equilibrium transport models. Lundstrom only actually considered the behavior of 226Ra and 231Pa in these models because he assumed that U/Th variations and 230Th variations reflect fundamental source differences (Lundstrom et al., 1995). Moreover, all the melts produced in this model are mixed at the top of the column according to flux and thus this model only produces a single-melt composition. Using this model, he found that 231Pa produced the strongest constraints on the suction parameter and required near equilibrium flow ðS , 0:05Þ: However, the depletion of rare-earth elements (REEs) in abyssal peridotites required near fractional melting with S , 0:5 – 0:8 depending on how depleted the solid source is assumed to be. In addition, because this model produces a single-melt composition at the top and was not used to explore 230Th variations, this version of the model was not used to address the local correlations on equiline diagrams. Nevertheless, it is possible to use this model to explore the behavior on an equiline diagram. Jull et al. (2002) developed a very similar model to that of Lundstrom (2000) and used it to try to model the observed correlations of (230Th/232Th) with (238U/232Th) on equiline diagrams and the anticorrelation between (226Ra/ 230 Th) and (230Th/238U) both seen at 98 N on the EPR. The principal difference between the Jull et al. (2002). and the Lundstrom (2002) approach is that Jull et al. (2002). use a more realistic partition coefficient structure that allows for garnet melting at the base of the model to produce significant 230 Th-excesses. More importantly, they do not mix the channel melts and the interchannel melts at the top of the model. Instead they assume that these melts form mixing end-members with the channel melts preserving high 230Th-excesses from the base of the melting zones (but having negligible (226Ra/230Th)), while the interchannel melts are high in 226Ra produced near the surface by porous flow at low porosities but have negligible 230Thexcesses. Jull et al. (2002) also allowed the depth where channels begin to be a variable and systematically explored a reasonably large region of parameter space in both physical parameters and partition coefficients. For a narrow parameter regime, Jull et al. (2002) were able to produce both (230Th/238U) and 226Ra variations that were comparable to observations for melts at 98 N EPR

U-series Melting Models (Sims et al., 2002) that had the same long-lived isotopic signature. For this regime, the channels were required to extend into the garnet stability field at depth and the calculation still required ,60% of the melt to flow through the interchannel regions. Thus, the solution is still principally an equilibrium porous flow model with some accelerated channeling. Jull et al. (2002) also attempt to explain the very depleted REE signatures of abyssal peridotites but have the same problem as Lundstrom (2002) in that the parameters required to match the U-series excesses in the model do not produce peridotites that are as depleted as observed. (However, as a small source of confusion, Jull et al. (2002) use the opposite convention for the suction parameter S. In their paper, S ¼ 1 is pure equilibrium flow and S ¼ 0 is fractional melting.) Nevertheless, the hybrid two-porosity models are promising and combine the important feature that transport models produce excesses in different parts of the melting regime, while channel models increase the transport rate of 230Th from depth. Nevertheless, the actual formation and structure of channels in these models is ad hoc as is the parametrization of “melt suction” S and the mechanics of mixing. While Jull et al. (2002) provide mixing end-members for their model, the model still only produces two melt compositions at any depth and the model does not include any specific mechanism for mixing. To go further requires consistent fluid-mechanical models that explicitly calculate the behavior of the melt channels. Preliminary work on such models (next section) suggests that the results of the simpler hybrid models will hold qualitatively and that U-series and stable element variability may finally be used to address the dynamics and structure of partially molten regimes. 3.14.4.3.5

Full reactive transport models

Motivated by observations of reactive dunite channels in ophiolites, models for reactive flow in deformable permeable media have been developed that show that fluid flow along a solubility gradient will lead to flow localization and the development of a coalescing network of highporosity melt channels surrounded by extremely low porosity regions (Aharonov et al., 1995; Spiegelman et al., 2001) (Figure 14). The key to this process is the inclusion of compaction. Because the channels are lower-pressure than the surrounding region, they do suck melt into them as was suggested by Iwamori (1993a) (although the suction is not coupled to the melting rate). The consequence of melt suction is to reduce the porosity of the interchannel region by compaction. If there is no resistance to compaction at small porosities, these models will actually compact to

501

Figure 14 (a) Porosity and (b) residual channel structure for the reactive flow problem described in Spiegelman et al. (2001). This solution assumes a static solid composed of an insoluble and soluble phase (e.g., ol and opx) where the solubility of the soluble phase increases upwards. This material is fluxed by a corrosive fluid that interacts with the solubility gradient and produces a localized network of high-porosity melt channels surrounded by compacted low porosity regions. In this run the soluble phase becomes exhausted from the channels leaving a residual network (dark regions) that qualitatively resembles residual dunites seen in ophiolites (e.g., Kelemen et al., 1995a, 2000; Braun and Kelemen, 2002).

near impermeability. The very small porosities in the interchannel region can cause significant chemical effects, particularly for U-series nuclides. Spiegelman and Kelemen (2003) explore the chemical consequences of these channel systems for stable trace elements and show that channel systems can produce significant diversity even from a homogeneous source in complete chemical equilibrium. The behavior is similar to that predicted from the two-porosity models. The interchannel regions become extremely depleted because the low permeabilities prevent deeper enriched melts from interacting with the solid. Meanwhile, the solid continues to melt, producing near fractional instantaneous melts in equilibrium with the solid (but extremely depleted). These melts are then sucked into the edges of the channel systems but not mixed. The end result is highly variable channel chemistry where the centers of the channels are deep, enriched melts that have mixed from across the entire column, while the edges are extremely depleted shallow melts. Thus, each channel delivers the entire range of chemical

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Melt Migration in Oceanic Crustal Production

variability to the crust on the scale of the channels (10 – 100 m). The melts produced show variability comparable to that seen in lavas and melt inclusions (Figure 16), while the residues (which are dominated by the interchannel regions) are extremely depleted and have compositions comparable to the abyssal peridotites (Figure 17). Figures 15 –17 show results from a high-resolution single-channel calculation showing the porosity structure and synthetic data sets sampled from the melts and residues that show good qualitative agreement with observations. The question remains, however, what these channel systems will do to U-series elements. Solving for U-series excesses directly in these systems is significantly harder to do accurately than in the one-dimensional column problems. However, we can use the same approach as in Spiegelman and Elliott (1993) to separate the effects of ingrowth from melting by assuming

that the total concentration of a nuclide can be written as cfi ¼ ai cfis

ð13Þ

where cfis is the concentration in the melt of a fictive “stable” element with the same partitioning behavior as the radioactive nuclide and ai is the “ingrowth factor.” We then solve for the log of the concentrations of U, Th, Ra, Pa as if they were stable elements using

›Uif þ veff · 7Uif ›t _i Gð1 2 Di Þ þ rs ð1 2 fÞD ¼2 rf f þ rs ð1 2 fÞDi

ð14Þ

where Uif ¼ lnðcfis =cf0 Þ is the log of the concentration of element i relative to the first instantaneous melt. Di is the bulk partition coefficient of element i, G is the total melting rate which includes contribution from both adiabatic and reactive melting, rf, rs are the melt and solid densities, f _ i ¼ ›Di =›t þ v · 7Di is the is the porosity, and D change in time of the partition coefficient in a frame following the solid: veff ¼

rf fv þ rs ð1 2 fÞDi v rf f þ rs ð1 2 fÞDi

ð15Þ

is the effective velocity of a tracer in chemical equilibrium with both melt and solid (e.g., see Spiegelman, 1996). Completely incompatible elements ðD ! 0Þ travel at the melt speed v; while completely compatible elements ðD ! 1Þ travel with the solid. The velocity field of intermediate elements depends principally on D=f; as usual. We then use the concentration ratios of the stable elements to solve for the excesses ðai Þ using

› ai þ veff · 7ai ¼ li ½Ri ai21 2 ai  ð16Þ ›t where Ri ¼ Figure 15 Porosity structure of a high-resolution single-channel calculation for an upwelling system undergoing melting by both adiabatic decompression and reactive flow (see Spiegelman and Kelemen, 2003). Colors show the porosity field at late times in the run where the porosity is quasi steady-state. The maximum porosity at the top of the column (red) is ,0.8% while the minimum porosity at the bottom (dark blue) is ,10 times smaller. Axis ticks are height and width relative to the overall height of the box. In the absence of channels this problem is identical to the equilibrium one-porosity transport model of Spiegelman and Elliott (1993). Introduction of channels, however, produces interesting new chemical effects similar to the two porosity models. See Spiegelman and Kelemen (2003) for details.

Di0 Dði21Þ0

!

f f rf f þ rs ð1 2 fÞDi21 Ui21 e 2Ui ð17Þ rf f þ rs ð1 2 fÞDi

controls the amount of ingrowth by weighting the importance of the parent nuclide. This term depends principally on the relative size of the porosity and partition coefficients as well as the partition coefficients at the first instance of melting and the ratio of concentrations of the “stable” form of the parent and daughter elements. While it looks somewhat complicated, this twostep process is much more accurate than solving for the concentrations directly and gives information that is readily plotted on an equiline diagram. For example, by this technique [230Th/232Th] is simply aTh because 232Th behave principally as a stable element for these problems.

U-series Melting Models

503

Figure 16 Comparison of model output and melt inclusion data from Spiegelman and Kelemen (2003). (a) Synthetic spidergram for highly incompatible elements plus rare-earths sampled from melts produced at the top of the melting calculation in Figure 15. The mean concentration is comparable to the maximum degree of melting but the distribution is similar to dynamic melting calculations. Unlike the single porosity models, these models produce a distribution of melt compositions at the top (rather than a single melt composition) even from a homogeneous source. The spidergrams here are sampled from this distribution in proportion to the melt flux (i.e., enriched melts from the center of the channel are more common than depleted melts). Note, this variability is produced within the channel on a small scale. (b) REE patterns from olivine hosted melt inclusions in three hand samples from the mid-Atlantic ridge (Sobolev and Shimizu, 1993, 1992; Shimizu and Grove, 1998). These samples show a similar range in variability. Most of the melt inclusions are similar in composition to the host glasses but there is a distinct tail of “ultra-depleted melts.”

Likewise, [238U/232Th] is calculated from the stable components. Thus, an equiline diagram simply plots the excess due to ingrowth against any chemical fractionation of parent and daughter. The equiline diagram in Figures 18(b) and 19(b) were constructed by this method. Figure 18 shows a preliminary calculation for U-series excesses in the single-channel problem shown in Figure 15 and is solved using the parameters in Table 3 with self-consistent bulk partition coefficients intended to span two endmember scenarios (Table 2). This problem has not been tuned to match the observations. Significantly, the bulk Ds are illustrative snap shots at a single pressure and compositions inappropriate for the whole melting column. Moreover, the channel is much larger than observed to improve numerical resolution. Nevertheless, as a first cut, this model can produce U-series distributions that are qualitatively similar to those observed. The important feature of these models that is different from most geochemical models, is that they produce a distribution of melt compositions at the top of the system, rather than a single-melt composition. These distributions can then be compared to spatially localized sets of samples. For example, Figure 18 shows the output of this model compared to U-series data from 98 500 N EPR (Sims et al., 2002). While the match is not ideal, the patterns are intriguing. This particular run produces a positive correlation on an equiline diagram even for a homogeneous

source. The fractionation of (238U/232Th) is solely due to melting with the melts from the center of the channels being enriched mixtures with deep sources and large 230Th-excesses, while the edges of the channels are depleted shallow melts with smaller excesses. A key prediction from these models is that there should be some correlation between U-series excesses and general trace element abundances for systems that can be demonstrated to have the same source. The partition coefficients used here are reasonably good to explain both 230Th-excesses as well as the range in (238U/232Th) although this run produces somewhat greater depletion than is seen in the data. Figure 18(c) shows a negative correlation between (226Ra/230Th) and (230Th /238U) although the model has lower excesses than observed. Finally, Figure 18(d) shows a positive excess between (231Pa/235U) and (230Th/238U) which does not appear to be present in the data. As with the single-porosity transport models, these calculations are sensitive to the parameters, particularly the partition coefficients. Figure 19 shows the same plots but for the bulk partition coefficients of Wood et al. (1999) from Table 2. As before, these Ds produce larger 226Ra- and 231 Pa-excesses but negligible 230Th-excesses. The models have not been explored yet to understand whether we might expect spreading rate dependence or not. These calculations suggest that much of the observed chemical variability of local ridge chemistry reflects the structure and

504

Melt Migration in Oceanic Crustal Production complexity. The previous section has attempted to clarify their pros and cons. We summarize here:

101 Channels

.

cs/c0s

100 Inter-channels

10–1

.

10–2 Chondrite normalized concentration

(a)

(b)

La Ce Nd Sm Eu Gd Tb Dy Ho Er Tm Yb 10

abyssal perid

1

0.1

b Cpx in Oman harzburgite

0.01 La Ce

Nd

Sm

Dy

OMB-43 OMB-31 OMB-30 OMB-27 OMB-23 OMB-22 OMB-16 OMB-14B OMB-14A

Er

Yb

Figure 17 Comparison of model output for distribution of residue compositions with abyssal peridotites and harzburgites from Oman. (a) Distribution of clinopyroxene compositions in equilibrium with melts at the top of the melting column corresponding to Figure 15. This distribution is sampled uniformly by area (as if we were sampling residues from the top of the column) and is dominated by depleted inter-channel samples. (b) Clinopyroxene compositions in harzburgites from the mantle section of the Oman ophiolite (Kelemen et al., 1995a) superimposed on the range of cpx compositions from abyssal peridotites (Johnson et al., 1990; Johnson and Dick, 1992), showing the predominance of highly depleted samples.

plumbing of the partially molten region. How spreading rate affects channeling is still undetermined. To go further clearly requires a systematic exploration of parameter space to understand the full sensitivity of the model and compare its outputs to both U-series and stable trace element data. However, the framework now exists and the data sets are just beginning to come together that will allow for future exploration of both theory and data.

3.14.5 SUMMARY OF MODEL BEHAVIOR The challenge of explaining observed U-series excess in MORBs has led to the development of a large number of models of differing levels of

.

.

.

.

Closed-system models cannot account for observed 230Th, 226Ra, and 231Pa-excesses for degrees of melting much larger than the bulk partition coefficients. As F , 10% for MORB while DU ; DTh ,, 0:01; these models are not viable for MORB generation. Both “dynamic melting” and “transport” models can produce the gross level of observed U-series excesses in MORB for appropriate parameters. However, they lead to different inferences on melt transport rates – Dynamic melting models produce all correlated excesses at the bottom of the melting regime and require melt transport rates sufficient to transport radium from the bottom of the melting regime on timescales short compared to the half-life of radium (1,600 yr). – Equilibrium transport models produce excesses throughout the melting column. In particular, thorium excesses are produced at depth in the presence of garnet and highpressure pyroxenes, while potentially observable radium excesses are produced near the top of the melting column. These models still require rapid melt transport near the top so that the porosity in the radium production zone is comparable to DTh. Both models assume a single porosity at any height in the melting region and produce well mixed melts at the top. Thus, these models cannot fractionate elemental uranium from thorium and cannot produce (230Th/232Th) versus (238U/232Th) correlations at constant 87 Sr/86Sr as is observed. Two-porosity models can begin to explain both excesses and correlations by allowing multiple melt compositions to be produced for the channel and interchannel regions. The actual mechanism and structure of channel formation is somewhat ad hoc but the underlying behavior seems justified qualitatively. Full fluid-mechanically consistent melt transport models with reactive channeling extend the results of the two-porosity models and produce distributions of compositions for both stable and radiogenic tracers in melts and residues. These models suggest that much information on the structure and rates of magmatic process might be contained in the observed variability of mantle melts but they need to be explored more rigorously. We have stressed accounting for the striking local variations evident in MORB U-series systematics. There are some less robust variations with geophysical parameters. Equilibrium melting one-porosity models

Concluding Remarks

505

Figure 18 (a) profiles of activity ratios as a function of distance across the top of the calculation shown in Figure 15 for (230Th/238U), (226Ra/230Th), and (231Pa/235U). Partition coefficients are after Salters et al. (2002), see Table 2. (b) U, Th behavior of this model on an equiline diagram. Blue squares show model results, gray circles are data from 98 500 N, EPR (Sims et al., 2002). These models actually produce correlated distributions on an equiline diagram because the channel melts tend to be more enriched deep melts with significant (230Th/238U) while the interchannel melts are shallow depleted melts with negligible 230Th-excesses. (c) Anticorrelation of (226Ra/230Th) and (230Th/238U) in these systems. Again, the interchannel melts have large Ra excesses while the channel melts have small ones. Ra excesses for this set of parameters are smaller than observed but show the observed anticorrelation. (d) Pa versus Th excesses. Pa is lower and shows a more positive correlation than observed.

appear to be able to reproduce the weak correlation of axial depth with 230Th-excess, and this is likely to extend to two-porosity models. The effect of spreading rate is still unclear observationally and the lack of clear correlation of absolute excesses potentially an interesting test of models.

3.14.6 CONCLUDING REMARKS U-series data place stringent constraints on melting and melt migration models and successful input parameters. Sophistication of models and analytical data have advanced considerably in the last decade. It is evident that a significant fraction of U-series variability is present on a local scale.

Detailed data sets that have a full range of geochemical tracers on the same samples analyzed for U-series disequilibrium will be particularly valuable for further developing the field (e.g. Sims et al., 2002). A limiting factor is finding sample suites clearly young enough that post-eruptive decay is insignificant (a problem most acute for 226 Ra – 230Th disequilibrium). Progress in developing shorter lived disequilibrium to screen for suitably young samples is thus very welcome (Rubin et al., 2001). Generating high precision U-series data sets is also time consuming. New plasma ionization multi-collector massspectrometers, however, should help improve both speed and precision of analysis. The results of recently developed melting models that can calculate both abundances and distributions of

506

Melt Migration in Oceanic Crustal Production Activities ratios at t = 85.00

Equiline diagram (238U/232Th)0 =1.20946, ([Th/U]=2.6)

8 2

9˚ N data model equiline

7 1.8 (230Th/232Th)

6 (230Th/238U) (226Ra/230Th) (221Pa/235U)

5 4

1.6 1.4

3 1.2

2 1

0

0.05 Distance x/h

(a)

1

7

1.4

1.6

1.8

2

Pa versus Th excesses

3.4

9˚ N data model

1.2

(238U/232Th)

(b)

Ra versus Th excesses

8

3.2 3

6

(221Pa/235U)

(226Ra/230Th)

1

0.1

5 4

2.8 2.6 2.4

3 2

2.2 1

(c)

1.1

2

1.2

(230Th/238U)

(d)

9˚ N data model

1

1.1

1.2

(230Th/238U)

Figure 19 The same figures as in Figure 18 but using the partition coefficients after Blundy and Wood (2003b) see Table 2. (a) horizontal profiles at surface; (b) an equiline diagram showing a small 238U-excess (but most samples are close to the equiline); (c) (226Ra/230Th) versus (230Th/238U); (d) (231Pa/235U) versus (230Th/238U). The large 226Ra- and 231Pa excesses are due to the considerably larger DU and DTh used in the upper part of this calculation.

Table 3 Parameters for full channel U-series model in Figure 18. Partition coefficients for these runs are taken from Table 2 with Salters et al. (2002) used for Figure 18 and McDade et al. (2003) (garnet) and Wood et al. (1999) (spinel) used in Figure 19. DRa ¼ DPa ¼ 1025 used in both figures. Variable Upwelling rate Column height Max. degree of melting w0 =W0 % box in Gt field

Value 7 cm yr21 80 km 10% 50 25%

chemical variability are encouraging. Further exploration of such models in conjunction with more comprehensive data sets holds promise for better understanding the processes of melt migration beneath ridges.

ACKNOWLEDGMENTS Discussions with many people over the past fifteen years have helped shape ideas in this contribution. More specifically, we are very grateful to Ken Rubin for a valuable review on a less-than-perfect initial version, which helped significantly improve the manuscript. Additional comments by Ken Sims, Chris Hawkesworth, and John Blundy were much appreciated. Discussions with Vincent Salters, Jon Blundy, and Bernie Wood were very useful in trying to synthesize partitioning data. We thank the editor for her patience and faith in our final delivery. The formative role of D. House in this endeavor is recognized. T. E. thanks the Leverhulme Trust for generous funding during the period of writing. NERC and NSF funding of T.E. and M.S., respectively, is also acknowledged.

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Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 465– 510

3.15 Hydrothermal Alteration Processes in the Oceanic Crust H. Staudigel University of California at San Diego, La Jolla, CA, USA 3.15.1 INTRODUCTION 3.15.2 THE UNALTERED OCEANIC-CRUST PROTOLITH 3.15.2.1 A “Standard Section” for the Oceanic Crust 3.15.2.2 Estimating Unaltered Oceanic Crust Compositions 3.15.3 DETERMINING THE ALTERED COMPOSITION OF THE OCEANIC CRUST 3.15.3.1 Recovery Rate 3.15.3.2 Types of Alteration 3.15.3.3 Duration of Alteration 3.15.3.4 Determining the Composition of Extremely Heterogeneous Altered Crust 3.15.4 CHEMICAL CHANGES IN ALTERED CRUST COMPOSITION DUE TO HYDROTHERMAL PROCESSES 3.15.4.1 Time Dependence of Crust Hydration and Carbonate Addition 3.15.4.2 Chemical Fluxes between Oceanic Crust and Seawater: Methods and Uncertainties 3.15.4.3 Chemical Fluxes 3.15.5 DISCUSSION 3.15.5.1 Hydrothermal Fluxes: Rock Data versus Fluid Data 3.15.5.1.1 Uncertainties 3.15.5.1.2 Bulk fluxes 3.15.5.1.3 Reconciling hydrothermal fluxes from fluid and rock data 3.15.5.2 Impact of Ocean-crust Composition on Arc Processes and Mantle Heterogeneity 3.15.6 CONCLUSIONS REFERENCES

511 512 512 515 517 517 518 519 520 522 522 523 524 529 529 530 530 531 531 532 533

hydrothermal fluxes also carry substantial elemental flux between seawater and the oceanic crust. Combined with ocean-crust generation and recycling, these processes produce a two-way geochemical pathway between the oceans and the mantle. Recycling of altered oceanic crust into the mantle is likely to produce some of the mantle’s chemical heterogeneity (e.g., Hofmann, 1988; see Chapter 2.04) and the delivery of mantle-derived materials to seawater through hydrothermal systems has profound effects on seawater chemistry (e.g., Wheat and Mottl, 2000; Chapters 3.15 and 6.07). Hydrothermal vents in mid-ocean ridges offer a unique habitat for very diverse biological communities that derive much of their energy needs from chemical energy in vent fluids (Jannasch and Mottl, 1985; Jannasch, 1995).

3.15.1 INTRODUCTION Hydrothermal alteration processes occurring in oceanic crust impact the physical, chemical, and biological processes of the Earth system. These hydrothermal systems are manifested in vents ranging from 350 8C black smokers, found exclusively in the axial zone of some ridge segments, to 20 8C low-temperature vents at the ridge axis or flanks. Collectively, these systems are responsible for ,20% of Earth’s total heat loss (11 TW; C. A. Stein and S. Stein (1994a,b)) and have major impact on ocean and solid earth chemistry. Elderfield and Schultz (1996) estimate black-smoker water fluxes to be ,3.5 £ 1012 kg yr21 and low-temperature fluxes to be ,6.4 £ 1014 kg yr21 (at 20 8C). These 511

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The interior of the oceanic crust is likely to host a deep-ocean biosphere that reaches to at least 500 m depth (Furnes and Staudigel, 1999). It is important to quantify hydrothermal chemical fluxes because they bear on the chemical and biological evolution of the Earth, the chemical composition of seawater, geochemical mass balance at arcs, and the heterogeneity of the mantle. Hydrothermal fluxes can be independently determined by analyzing the composition of hydrothermal fluids or by analyzing the alterationrelated chemical changes in the oceanic crust. Ideally these two methods should yield the same results, but a comparison of data shows that there are major discrepancies between these types of estimates (e.g., Hart and Staudigel, 1982; Chapter 3.15). Reconciling these discrepancies is important for improving our understanding of this central theme in Earth system sciences. This review focuses on chemical flux estimates derived from studies of the oceanic crust, exploring in detail how such estimates are made, and the underlying assumptions and uncertainties. Three main themes will be covered. The first focuses the role of the original igneous characteristics of the crust in determining the nature of hydrothermal alteration processes. This includes how primary lithology and composition influence alteration, and difficulties encountered in determining an unaltered “fresh-rock” baseline composition for any particular ocean-crust section. The second theme focuses on the methods by which the bulkaltered oceanic composition is determined, and the attendant uncertainties. These include the difficulty of determining an average composition of a very heterogeneous medium by the analyses of rather small samples, and the limitations imposed by an incomplete sampling process on the ocean floor. Finally, hydrothermal fluxes inferred from ocean-crust data are compared to fluxes from hydrothermal vent studies and the reasons behind their differences are explored.

3.15.2 THE UNALTERED OCEANIC-CRUST PROTOLITH Determining the original igneous characteristics of altered oceanic crust is an important step in understanding its alteration. There are two very different reasons for this. First, the primary igneous characteristics of the oceanic crust determine important aspects of alteration by controlling heat exchange, fluid flow, and the reactivity of crustal materials with hydrous fluids. These primary characteristics may vary as a function of spreading rate, ridge morphology, or position in a volcano, producing a range of alteration types. Effectively integrating these different crust types is an important

challenge in producing a meaningful global integration of alteration fluxes. This problem is addressed here by defining a “standard” section for normal oceanic crust based on results from available sample materials. Second, the original composition of oceanic crust plays an important role in determining chemical fluxes. Uncertainties in the original composition translate directly into uncertainties of chemical fluxes, and in many cases these differences are the main sources of errors. Both of these features are addressed in the following sections. 3.15.2.1 A “Standard Section” for the Oceanic Crust The most comprehensive understanding of processes occurring in the oceanic crust comes from the study of subaerially exposed ocean-crust sections in ophiolites, which offer extensive and often continuous exposure over very large areas. The observations from ophiolites complement the significantly more limited data from submerged, in situ oceanic crust, which in turn serves as an important test for the validity of these subaerial analogues. For this reason, ophiolites are key to understanding the relationships between alteration behavior and primary crust characteristics. The main ophiolites and ocean-crust drill holes discussed in this paper include the Troodos ophiolite in Cyprus, the Samail ophiolite in Oman, and ocean-crust drilling sites 332B and 417/418 in the Atlantic Ocean and sites 504B and 801C in the eastern and western Pacific, respectively (Figure 1, Table 1). The classic view of the oceanic crust is based on the “Penrose” ophiolite assemblage (Penrose Conference Participants, 1972) that includes a characteristic sequence of rock types: basal mafic cumulates are overlain by gabbro-norites, gabbros, sheeted dikes, and pillow lavas on the top. Pillow lavas may be overlain by hydrothermal metalliferous sediments, that are commonly overlain by pelagic, typically siliceous fine-grained sediments (cherts). This three-decade-old ophiolite model of the oceanic crust remains effectively unchanged in terms of the main rock types found in oceanic crust and in terms of their stratigraphic succession. The main modifications to this model include changes in the average thicknesses of units. In addition, there is some in situ oceanic crust that appears to lack any extrusives, but this type of ocean crust is mostly confined to very slow spreading ridges and represents probably ,5% of the total crust produced (e.g., Bach et al., 2001). Study of the global impact of ocean-crust generation and alteration ideally requires integration of data from all major crustal types, if possible, including several studies that offer

The Unaltered Oceanic-crust Protolith

513

Troodos ophiolite Samail ophiolite

332B 417A,D 418A 543A

801C

896C 504B 765C 735B

300

Figure 1

0

60

120

180

Locations of the drill sites and ophiolites discussed in the text.

Table 1 Ophiolite and ocean-crust drill hole locations. Site

Lat.

Lon.

Age (Ma)

Ocean/Region

Depth (m)

Recovery (%)

Comments

580 206

21 70

366 544 2,000 44 1,000 247 135

70 72 20 81 31 60

290

28

W MAR Highly altered 150 m abyssal hill Normal crust Normal crust MORB extrusives/dikes MORB extrusives MORB gabbros/troctolites Normal MORB extrusives Alkalic basalts and normal MORB extrusives Normal MORB extrusives

332B 417A

36.945 25.106

33.641 268.041

3.5 120

Atlantic Atlantic

417D 418A 504B 543A 735B 765C 801C

25.106 25.035 1.227 15.712 232.722 215.976 18.642

268.041 268.057 283.734 258.654 57.264 117.592 156.360

120 120 5.9 82 9 155 157

Atlantic Atlantic E. Pacific Atlantic Indian Ocean Indian Ocean W. Pacific

896C Troodos CY 1

1.217 34.800 35.048

283.723 32.000 33.179

5.9 92

Tethys/Cyprus

1,150

95

Suprasubduction zone extrusives

CY 1A CY 4 Samail

Tethys Tehtys/Oman

95

58.000

92 120

2,300

24.500

Dikes gabbros, ultramafics Normal MORB

E. Pacific

All locations given in decimal degrees, negative values give southern latitudes and western longitudes.

independent constraints. Unfortunately, there are too few sections studied for such coverage, and for this reason one has to combine all observations into one “general” type of ocean-crust section. This is clearly an over-simplification, but to a first order a good approximation for the most abundantly produced oceanic crust. Using the stratigraphy of the classic Penrose ophiolite and more modern estimates of unit thicknesses, our standard section (Figure 2) has a total crustal thickness of 7.1 km (White et al., 1992; C. Z. Mutter and J. C. Mutter, 1993), which includes 1,000 m of extrusives, 1,100 m of sheeted dikes, and 5,100 m of gabbros. This “standard” section translates into a volume production rate of 24.14 km3 yr21 using Parsons’ (1984) estimate of the ocean-crust surface area

production rate of 3.4 km2 yr21. This rate of oceanic-crust production can be translated into a mass production rate by using in situ densities measured in the oceanic crust. These values range from ,2,700 kg m23 for extrusives (Salisbury et al., 1979), to ,3,000 kg m23 for gabbros (Dick et al., 2000). The sheeted dikes are extrapolated to have an average density of 2,850 kg m23. This yields a total ocean-crust production rate of 7.05 £ 1016 g yr21. However, it must be emphasized that there are still significant uncertainties in these total fluxes and production rates. The Troodos and Samail ophiolites (Figure 1) contain all the essential components of our standard section in Figure 2. Both of these

514

Top

1,000

Hydrothermal Alteration Processes in the Oceanic Crust CY 1 418A 504B CY1A

Extreme alteration:

Extrusives 1,000 m / 2,700 kg m–3 9.18 × 1015 g yr–1

Upper extrusives (600 m)

Gossans massive sulfides

Lower extrusives (400 m)

Spilites

Sheeted dikes 2,000

1,100 m / 2,850kg m–3 1.06 × 1016 g yr–1

Sheeted dikes (1,000 m)

3,000

Gabbros ultramafics 4,000 5,000m /3,000 kg m–3 5.1×1016 g yr–1

CY4 Epidosites 735B Plagiogranites, amphibolites

Gabbros (3,250 m)

5,000 Gabbro-norites (1,000 m)

6,000

Serpentinites, chromitites, talcum

Ultramafics (750 m) 7,000

Total 7,100 m 7.08 × 1016 g yr–1

Figure 2 The ocean crust reference section used in this paper, using a standard “Penrose” style ophiolite assemblage (Penrose Conference Participants, 1972) and a crustal thickness after White et al. (1992) and C. Z. Mutter and J. C. Mutter (1993). Various modeling parameters used are also indicated, including densities and fluxes, the approximate positions of drill holes discussed in this paper, and some typical extreme alteration environments found at various depths of the oceanic crust.

ophiolites were formed in the ancient Tethys Ocean (92 Ma for the Troodos Ophiolite and 100– 130 Ma for the Samail ophiolite). The Samail ophiolite represents crust that is quite similar to “normal” oceanic crust (Kelemen et al., 1997), but it has not been studied for geochemical mass balances. The Troodos was studied extensively for its primary and hydrothermal characteristics but it reflects an oceanic crust that is more typical of supra-subduction zone settings, having geochemical characteristics distinct from normal mid-ocean ridge basalt (MORB) (Robinson et al., 1983; Schmincke et al., 1983). For this reason, the bulk composition of the Troodos ophiolite is not typical of average oceanic crust, even though it is clearly a part of the spectrum that is produced and recycled. The Troodos ophiolite is well studied, partly because of an intense drilling campaign that yielded a total of 8,000 m of core material from all structural levels in the crust. Key drill sites include CY1 and CY1A, with a total penetration of 1,175 m though extrusives (Gibson et al., 1991a), and CY 4, that penetrated 2.3 km of sheeted dikes, gabbros, and pyroxenites

(Gibson et al., 1991b). Their stratigraphic position is plotted in the reference section in Figure 2. Recovery from all of these drill cores was found to be ,100%. There are a large number of drill sites in oceanic crust, but four sites stand out for their impact on studies of hydrothermal alteration processes. These include sites 417 and 418, which were studied to constrain the chemical mass fluxes of the upper crust, site 504B for the mid crust or sheeted dikes and site 735B for the gabbro section (Figures 1 and 2; Table 1). Three drill holes at sites 417 and 418 probe upper oceanic crust of the western Atlantic. This crust formed at the relatively slowly spreading mid-Atlantic ridge, which may be more fractured than fast spreading crust, and contain more pillows and less massive flows. Site 417A, was drilled into a 150 m high abyssal hill with a total penetration of 208 m, whereas sites 417D and 418A were drilled in the nearby abyssal plain, with penetrations of 360 m and 544 m, respectively. Recovery rate was about 70% at all of the sites. Site 417A contains, on average, 20%

The Unaltered Oceanic-crust Protolith volcaniclastics, whereas sites 417D and 418A contain only ,6% volcaniclastics (Robinson et al., 1979). The higher abundance of volcaniclastics and the more exposed nature of Site 417A resulted in more intense alteration than at sites 417D and 418A (Donnelly et al., 1979a,b,c). Lithologies at Sites 417D and 418A are quite similar, but the crust was formed at different mid-ocean ridge volcanoes, offering a comparison between two separate but quite similar crustal sections. Staudigel et al. (1995, 1996) combined these sites into one representative section of almost 600 m length, including 20% materials from 417A and 40% each from 417D and 418A in the upper 150 m, all adjusted to 6% volcaniclastics. Additional sites that sample upper oceanic crust include 332B, 504B, and 801C (Figure 1; Table 1). Site 332B was also drilled in the Atlantic on very young crust, and 504B and 801C were drilled in the eastern and western Pacific, respectively. Amongst those, site 504B is the deepest drill hole into ocean crust, sampling both extrusives and the underlying sheeted dikes, thus offering a unique opportunity to study alteration processes in the deeper reaches of the crust (Alt et al., 1993a,b). Site 801C is located in the western Pacific, on 170 Myr old oceanic crust and penetrated almost 500 m of basalt with average recovery rate of 47% (Plank et al., 2000). This hole recovered alkalic basalts near the top, overlying a substantial section of “normal” oceanic crust below. The deeper portion of this hole is still under investigation. Samples of the deepest oceanic crust are accessible in only one ocean drill core: site 735B was drilled on 11 Myr old crust on the SW Indian Ridge (Dick et al., 2000). This hole penetrated a few tens of meters of pillow lavas and ,1,200 m of gabbros, with ,100% recovery. The site never reached cumulates (Dick et al., 2000). The lack of sheeted dikes and the near absence of pillow lavas is clearly quite different from the “normal” ophiolitic crust, but site 735 rock types are not very different from materials found in the upper plutonic section of ophiolites. Thus, site 735B is used to represent the deeper crust for the composite crust section. 3.15.2.2 Estimating Unaltered Oceanic Crust Compositions Estimating the unaltered composition of the oceanic crust is necessary for flux estimates because these estimates are determined as the difference between original and altered compositions. Table 2 displays an average unaltered mid-ocean ridge basalt composition (MORB, Hofmann, 1988) and estimate of an unfractionated MORB composition (Sun and McDonough, 1989), but it is important to emphasize that the

515

composition of MORB can vary substantially (e.g., Chapter 3.13 or http://petdb.ldeo.columbia. edu; Lehnert et al., 2000). This variation is too large for most elements to allow the use of a single fresh “average MORB” as a starting composition in chemical flux estimates. Thus, fresh compositions have to be constrained for any particular suite of altered basalts studied. There are two main ways in which the “freshrock” composition of an altered rock is determined. One is a petrographic – geochemical approach, which involves re-assembling a rock from the chemical analyses of residual igneous phases still present in the rock using their original modal abundances. Another is a purely geochemical approach that uses the chemical systematics in fresh rocks and placement of an altered rock in such fresh-rock systematics using alteration insensitive parameters. Deviations from the extrapolated fresh-rock composition are then interpreted as being caused by alteration processes. In both methods, there are significant errors. Interpretations of chemical trends have to contend with the natural scatter of sample distributions in correlation diagrams, as well as the cumulative analytical errors introduced from regressions, normalizations, and analytical uncertainty. Petrographic – geochemical reconstructions are affected by small-scale heterogeneities in mineral distribution and uncertainties in modal estimates. These errors are especially large for major elements or trace elements that are compatible in olivine and plagioclase, including magnesium, calcium, silicon, sodium, aluminum, iron, nickel, and barium. These errors are the lowest for incompatible elements that are present in very low abundances in the original igneous rock, such as H2O, CO2, uranium, potassium, rubidium, caesium, REEs, and thorium. A third approach that is commonly used to constrain chemical fluxes compares differently altered materials, such as altered pillow margins and less altered pillow interiors, or samples with or without alteration haloes around veins (e.g., Alt et al., 1986), mineralized and unmineralized zones or differently altered gabbros (e.g., Bach et al., 2001) in order to constrain chemical changes associated with alteration. However, “least” altered samples only rarely reflect the original composition reliably. A second problem in this approach is the relatively small sample sizes typically analyzed from ocean drilling materials. Typical sample sizes are about 15 cm3, which is small when compared with local variability in modal mineralogy. Indeed, individual phenocryst phases can be several millimeters in size. Local variability in modal mineralogy is particularly common in pillow lavas where phenocryst abundances can vary as a function of radial distance from the center or vertically within the center

Table 2 Fresh MORB

SiO2 Al2O3 FeOtot MnO MgO CaO Na2O K2 O Rb Cs CO2 H2O S Li B Sr U

50.45 15.26 10.43 0.19 7.58 11.3 2.679 0.11 1.26 0.0141 0.15 0.2 960 4.5 0.5 113 0.0711

Units

wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% ppm ppm wt.% wt.% ppm ppm ppm ppm ppm

Bulk rock gains (þ) and losses (2) for extrusives 0–600 m

Error

1.18 1.94 20.01 0 20.19 2.14 0.15 0.54 10.3 0.183 3.26 2.81 20.06 2.8 25.7 22 0.3

1.5 3 0.03 0.01 1 1.5 0.2 0.15 3 0.050 0.5 1

10 5 0.05

Dikes

Gabbros

Hydrothermal fluxes.

Average crustal gains/losses

Units

Total flux from crust

Units

600 –1000

Hydrothermal fluxes from submarine vents (g yr21) Flank

0.5

0.25

0.75 0 0.3 0.15 0.05 1 0.03 0.5 2.09 20.06 2.8 4 3 0.1

20.035 0 0.04 0.15 0 0 0 0.06 1.5 20.06 20.5 5.6 0.3 0.01

0.1

20.2 20.1 0.03 0 0 0 0.06 0.11 0.02 22 2.2 21 0.007 0.3

0.23704225

20.15690141 0.13352113 0.06549296 0.0484507 0.92676056 0.01715493 0.35521127 0.44873239 20.0037 22 4.81 1.4 0.03746479

g/100g

g/100 g g/100 g g/100 g g/100 g mg/kg mg/kg g/100 g g/100 g g/100 g mg/kg mg/kg mg/kg mg/kg

21.67E þ 14 29.65E þ 13

g yr21 g yr21

1.11E þ 14 29.43E þ 13 24.63E þ 13 21.71E þ 13 26.55E þ 09 21.21E þ 08 22.51E þ 14 23.17E þ 14 2.59E þ 12 7.71E þ 09 23.40E þ 11 29.68E þ 10 22.65E þ 09

g yr21 g yr21 g yr21 g yr21 g yr21 g yr21 g yr21 g yr21 g yr21 g yr21 g yr21 g yr21 g yr21

23.60E þ 14

g yr21

River fluxes (g yr21)

Axis low

Axis high

1.08E þ 12

1.80E þ 13

6.61E þ 13

3.85E þ 14

0.00E þ 00 3.69E þ 10 22.18E þ 14 2.64E þ 14 23.84E þ 13 23.11E þ 13 22.22E þ 09

1.65E þ 12 4.97E þ 11 24.03E þ 13 3.03E þ 11

1.37E þ 13 3.97E þ 12 21.05E þ 14 1.23E þ 14

6.12E þ 12 1.37E þ 10

3.01E þ 13 2.39E þ 10

27.04E þ 12

21.76E þ 12

27.04E þ 12

1.65E þ 12 3.48E þ 11 2.18E þ 14 6.73E þ 14 3.53E þ 14 1.22E þ 14 3.16E þ 10 6.40E þ 08 1.41E þ 15

23.53E þ 13 21.25E þ 10 1.84E þ 11 2.19E þ 11

21.60E þ 13 4.8587E þ 10 6.49E þ 09 29.00E þ 08

24.49E þ 13 4.37E þ 11 7.89E þ 10 4.60E þ 09

2.85E þ 13 9.72E þ 10 5.84E þ 11 2.02E þ 11 9.60E þ 09

1.90E þ 14

21.14E þ 13

1.02E þ 14

2.23E þ 15

Determining the Altered Composition of the Oceanic Crust of a pillow. Staudigel et al. (1996) estimated errors in chemical composition on the basis of variation in phenocryst modal abundances and found relatively large errors for some elements, particularly MgO (e.g., 7.7 ^ 0.2 wt.%). One of the key objectives in estimating fresh rock compositions is to determine the chemical inventory present before alteration in a given reference volume of altered rock. This is different from the “fresh-rock composition,” because water – rock interaction occurs in an open system. The pitfalls arising from open-system behavior can be illustrated in two examples: secondary mineral precipitation and rock dissolution. In rocks with significant pore space, minerals may precipitate in these pore spaces during alteration. For example, precipitation of 5 g per 100 g of calcium carbonate in the pore spaces of a fresh rock dilutes the abundances of all other chemical components by 5%. In this example, carbonate precipitation causes an apparent loss of ,2.5 wt.% SiO2, when in fact SiO2 remained immobile during the alteration. The opposite effect may be caused by wholesale (congruent) dissolution of the rock whereby massive amounts of basalt could be lost from a given rock volume, without any trace of concentration change. However, truly congruent dissolution is quite unlikely; dissolution typically leaves behind some residual immobile material. Titanium is generally assumed to be immobile, and the abundance of titanium in a rock can therefore be used to quantify open-system chemical behavior. The first step in addressing open-system behavior is to estimate the titanium concentration in the fresh rock equivalent, which is taken as the mass of titanium originally present in a rock per 100 g of the material analyzed. The fraction of titanium originally present, together with an independent estimate of the fresh rock composition, allows an estimation of the concentration of all the other elements in the fresh rock. Chemical fluxes, then, are calculated as the differences between elemental concentration in the unaltered rock (per 100 g analyzed) and the concentration of the element in the altered material. The most important sources of uncertainties in this procedure arise from the uncertainties in titanium determination, errors in the estimate of the titanium concentration in the unaltered rock (e.g., Staudigel et al., 1995, 1996) and errors due to small-scale titanium variations when using the titanium from a “fresh”-altered sample pair (e.g., Bach et al., 2001). It is possible to estimate the validity of the constant titanium assumption through a comparison with other, independent methods, such as working with constant reference volumes and estimating the amount of initial rock present based on density and pore space consideration.

517

Specifically, if the density is known for the fresh and the altered rocks, fresh and altered composition can be expressed on a volume percent basis, and fluxes are calculated relative to a constant volume. This method was evaluated at sites 417 and 418 and shown to produce almost identical results as the constant titanium assumption (Staudigel et al., 1996). 3.15.3 DETERMINING THE ALTERED COMPOSITION OF THE OCEANIC CRUST Determining the composition of altered oceanic crust is also not very straightforward. Several key steps are required in obtaining a meaningful estimate of altered rock compositions. First, the choice of a study site is important. Is the section old enough to have experienced the bulk of its alteration processes? Are the alteration patterns representative of most oceanic crust, or are they only of local importance? Is core recovery sufficient to allow a representative estimate of the section drilled? Once these three questions are answered positively, a method has to be worked out to determine the bulk composition of the crust and chemical fluxes on scale lengths that are meaningful for global chemical budgets. The following sections evaluate problems related to these questions, in particular the role of recovery, the range of “typical” types of alteration, constraints on the duration of alteration in the crust and techniques for determining representative compositions of a heterogeneous medium. These problems are illustrated with examples from particular ocean drill sites or ophiolites. 3.15.3.1 Recovery Rate It is extremely rare in ocean drilling to have 100% recovery of basement materials. Most drill cores from basement consist of variously rounded fragments that typically don’t match up with the neighboring core fragments. Recovery rates are low, particularly in the upper and young oceanic crust, which has not been effectively sealed through mineral precipitation and fractured on the same scale length as the drill bit (and smaller). Competent rock fragments, resistant to the cutting action of a drill bit, are held together by relatively soft materials like clays, carbonates, or chlorite. The stress imposed on resistant rock fragment is likely to break up the loosely cemented formation, and grind up the soft material filling the void and fracture spaces in between. This material is then ejected from the hole with the drilling fluid on the ocean floor as mud, sand, or chips and not in the core barrel. This problem also affects the inner walls of the drill hole, where competent rocks are

518

Hydrothermal Alteration Processes in the Oceanic Crust

more likely to be exposed than soft rocks. Most alteration materials are not found in distinct horizontal layers, but occur as pockets between or within pillows, or as veins that commonly cut the hole at very steep angles, which can be completely eroded, even though the neighboring material is actually recovered. A similar sampling bias occurs in outcrops of ophiolite sections, where clays are much more likely to be eroded than fresh, well-cemented basalt. These problems can be overcome only by complete recovery of drill core. The recovery rates of various ocean-crust drill holes are given in Table 1. Recovery rates are determined as the ratio of the cumulative length of recovered fragments to the actual length of the section drilled. This estimate is quite meaningful at high recovery rates, but not at low recovery rates, where individual pieces tend to be rounded off and have a smaller diameter, resulting in volume recovery rates that are substantially smaller than the linear rates. Thus, low recovery rates practically eliminate substantial portions of a drill hole from direct study and they can make it very difficult, if not impossible, to estimate bulk compositions from bulk rock data, particularly in pillow sections and breccia zones. Drilling in oceanic crust often has very low recovery, on the average of 25– 30%, with some sections having nearly zero recovery (e.g., a substantial fraction of 504B). At an average recovery rate of 30%, proportionately more material is recovered from massive units (50 – 100%), than breccia zones (0 – 10%) complicating the estimates of alteration inventories. Ocean drill sites with the highest recovery rates include site 801C, with almost 50% recovery, 417A, 417D, and 418A, with over 70% recovery, and 735, with 86% recovery. The Cyprus drill cores CY 1 and CY 4 have nearly 100% recovery. The other deep drill sites discussed here, 504B and 332B have substantially lower recovery rates, on the order of ,20 –30%. 3.15.3.2 Types of Alteration One of the key goals in quantifying alteration is to determine the respective contributions of

different types of alteration. The focus of these efforts have to be on what is considered “normal” or “most representative,” but it is also important to explore the complete range of alteration behavior in the oceanic crust. Even unusual alteration environments may have a significant impact on the geochemical behavior of some elements, and they may cause distinct geochemical behavior in subduction zones or in the mantle. For this reason, it is important to focus on the “average” geochemical behavior, but keep an eye on the compositional diversity as well. Studies of ophiolites have been particularly useful in identifying types of alteration in oceanic crust and for understanding their relative significance. Many of these alteration features have also been found in ocean drill cores, but many aspects of seafloor hydrothermal alteration remain unexplored by ocean drilling. The upper extrusive oceanic crust (0 – 600 m, Figure 2) is primarily altered at low temperatures (,100 8C). Alteration is commonly not pervasive, whereby igneous phases (glass, phenocrysts) may coexist with alteration phases (clays, zeolites, and carbonates). At high water – rock ratios, oxidative mineral assemblages (e.g., celadonitic clays) form intensely altered zones, but most ocean-crust alteration is accomplished by more reducing fluids at lower fluid –rock ratios (Table 3). However, in most low-temperature altered basalts, highly reducing minerals (like pyrite) may coexist in close proximity with oxidizing minerals like celadonite or hematite. The distinctive differences between mostly oxidized and reduced alteration environments are well illustrated by comparing the oxidative upper (abyssal-hill) portion of site 417A to the more reduced sections at sites 417D and 418A (e.g., Donnelly et al., 1979a,b,c; Alt and Honnorez, 1984). A similar contrast exists between Cyprus drill cores CY 1 and CY 1A, whereby CY 1, from the uppermost Akaki canyon, closely resembles the oxidized alteration in site 417A and the remaining core shows alteration behavior more like site 417D and 418A (Gillis and Robinson, 1988; Bednarz and Schmincke, 1989). Based on the distribution of alteration features in the northern portion of the Troodos ophiolite, Staudigel et al. (1995)

Table 3 Secondary minerals in the oceanic crust. Upper extrusive crust Oxidizing conditions Aragonite, analcite, calcite, celadonite, chalcedony Fe-hydroxide, hematite, philipiste, K-feldspar, saponite

Deeper extrusives Nonoxidizing

Anhydrite, analcite, calcite, celadonite, Fe-hydroxide, mixed layer chlorite – smectite, Na zeolite pyrite, saponite

Albite, calcite, chlorite, epidote, pumpellyite, prehnite, quartz, sphene

Determining the Altered Composition of the Oceanic Crust suggested that the 417A style of oxidative alteration makes up ,20% (^10%) of the upper 150 m of oceanic crust, while the reduced style of alteration seen at sites 417D and 418A are representative of 80%. The upper oceanic crust can show extreme compositional variation near hydrothermal vents. Based on the abundance of hydrothermal deposits in the Troodos ophiolite, and the frequency of black-smoker-type deposits in mid-ocean ridges, these highly mineralized zones are likely to represent less than one-tenth of a percent of the volume of the total upper oceanic crust. Thus for most elements, these deposits are not very important in the total mass balance. The exception to this is elements that are enriched in vent deposits by several orders of magnitude relative to basalt (e.g., copper, zinc, lead, manganese, iron, nickel, cobalt, platinum, silver, gold; see also Chapter 3.12). This limits the significance of these deposits to trace-element mass balances in the oceanic crust. A convincing case for this was made by Peuker-Ehrenbrink et al. (1994), who suggested that global lead cycles in seawater might be substantially influenced by the precipitation of a small amount of hydrothermal metalliferous sediments on top of the oceanic crust. In the deeper extrusive oceanic crust and in the sheeted dikes, alteration temperatures increase (.100 8C), water –rock ratios decrease, and most primary igneous phases tend to be almost entirely replaced by secondary phases (the exception again is near hydrothermal conduits, where temperature and water – rock ratios are high). In this depth range glass, olivine, and calcic plagioclase are typically replaced by greenschist-facies mineral assemblages (Table 3). Typical phase assemblages in deep extrusives and dikes for in situ oceanic crust are best described at site 504B (Alt et al., 1986, 1993, 1996), offering important insights into alteration processes in this depth range. Several important alteration environments, however, are not observed for in situ oceanic crust. These include spilites (Cann, 1969), a rock type that displays almost complete exchange of calcium for sodium, leading to formation of an albite-rich rock, and epidosites, the metal depleted epidote –quartz – chlorite assemblage that is likely to be characteristic of the reaction zones, or at least last equilibration zones of black-smoker fluids (e.g., Schiffman and Smith, 1988; Richardson et al., 1987; Bettison-Varga et al., 1992). In particular, the importance of spilites in chemical mass balances could be large, but they are effectively unknown for in situ oceanic crust. Spilites are therefore not considered in any mass balances. Epidosites are likely to be important for the mass balance of some trace metals in the

519

oceans, but are unlikely to influence the ocean crust generation – subduction budget because they are probably just as uncommon as the massive sulfide deposits. Another potentially important alteration assemblage is predicted from experimental studies and observations from black-smoker chemical compositions. Heating seawater to 450 8C results in massive precipitation of sulfate (mostly anhydrite; Bischoff and Seyfried, 1978) and may result in boiling, leaving behind brines (Butterfield et al., 1990). Anhydrite and brines are likely to be stored only temporarily in the oceanic crust and will be dissolved by seawater circulating through the crust at a later time after much of the magmatic heat is exchanged with seawater (e.g., Alt, 1994, 1995). Deeper levels in the oceanic crust display higher-grade hydrothermal alteration, ranging from amphibolite grade to anatexis (Table 3), where the formation of plagiogranites has been associated with the partial ingestion of hydrothermally altered ocean crust materials. Hydrothermal alteration processes in this regime have been studied in materials recovered from oceanic fracture zones, particularly in site 735B (Dick et al., 2000), as well as in ophiolites (Nehlig et al., 1994). Overall, it appears that the upper 200 m section of site 735B displays unusually high water –rock ratios due to its proximity to seawater, but the bulk of this 1.5 km section are more typical of normal crust (Dick et al., 2000). However, 735B does not display any of the extreme varieties of alteration observed in some ophiolites or in near-surface exposures on the ocean floor, such as serpentinites, or other hydrous equilibrium assemblages such as talcum deposits (e.g., Bonatti, 1976; Baer, 1963). Such compositional domains are likely to be formed at relatively slow spreading centers where brittle deformation may carry water to great depth, deeper than in the faster-spreading oceanic crust. Such deposits could potentially be very important for mass balances, particularly for recycling of water, but are not included in this review due to lack of data. 3.15.3.3 Duration of Alteration Determining that alteration is complete or near complete is one of the major pre-requisites for determining alteration fluxes at any study site, and for this reason, it is important to determine the duration of alteration in the oceanic crust. There are two main constraints for this: measurements of heat flow, which indicate the duration of convective heat loss, and isotopic dating of secondary minerals that precipitate during alteration of the oceanic crust. Heat-flow measurements can be used to determine the duration of seafloor alteration by

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Hydrothermal Alteration Processes in the Oceanic Crust

comparing the absolute estimate of heat flow to heat flow calculated from plate-cooling models. The difference between the two is then assigned to convective heat flow. The difference between theoretical and measured heat flow suggests total convective heat loss of approximately 11 TW (Stein et al., 1995). About one-third of this heat loss occurs within the first million years after crust formation, the second one-third occurs between 1 Ma and 8 Ma, and the last one-third occurs between 8 Ma and 65 Ma (Stein et al., 1995; Figure 3). However, the “termination” of convective heat loss at 65 Ma also coincides with the time beyond which the simple relationship between ocean-floor depth and plate-cooling models break down. Thus, a comparison of measured and theoretical heat flow is not very meaningful at this age. Alternatively, the duration of hydrothermal convection in the oceanic crust can be estimated by mapping the distribution of nonlinear temperature profiles taken during heat-flow measurements as a function of oceanic-crustal age. Purely conductive heat loss (i.e., no hydrothermal circulation) results in linear temperature profiles in sediments, while convective heat loss results in concave or convex profiles, depending on whether the water penetrates into or comes out of the sediments. Both heat-flow methods deliver similar results, suggesting that convective heat loss, hence hydrothermal alteration, may occur for time periods 100

of up to ,65 Ma (Parsons and Sclater, 1977; Langseth et al., 1988; Anderson et al., 1979; Stein et al., 1995; Pelayo et al., 1994). The duration of chemical exchange between seawater and basalt can also be determined by isotopic studies of hydrothermal minerals. A variety of techniques have been used, ranging from direct dating by K/Ar and Rb/Sr isochron techniques to comparisons of the initial strontium isotopic composition of alteration minerals with the isotopic evolution of seawater (Gallahan and Duncan, 1994; Richardson et al., 1980). Compiled isochron data for the duration of alteration at a series of DSDP and ODP sites are shown in Figure 3. The ages thus calculated are the differences between crustal age and veinmineral age. Due to the rather large spread of dates, ages were binned into 20 Ma intervals, with averages at the mid point of each interval shown in Figure 3. Even though rather crude, this distribution shows a remarkable resemblance to the heat-loss curve. While the resolution of sampling does not allow any strong constraints on the actual age distribution in the (most critical) first 20 Ma, they do show very clearly that vein-mineral precipitation can continue for a rather long time in this global distribution of sampling sites. Collectively, these observations suggest that oceanic crust younger than 10 Ma is unlikely to have experienced the complete cycle of seafloor alteration.

3.15.3.4 Determining the Composition of Extremely Heterogeneous Altered Crust

80 % vein mineral ages 60 40 Fraction of total convective heat loss 20

0

20

40 Million years

60

80

Figure 3 Cumulative convective heat loss (Booij et al., 1995) of the oceanic crust and completion of vein mineral deposition (Stein et al., 1995) in the oceanic crust. Heat flow curve after C. A. Stein and S. Stein (1994b), including vein mineral ages from site 261 after Hart and Staudigel (1986) site 417/418 after Hart and Staudigel (1978, 1986) and Richardson et al. (1980), site 462A after Hart and Staudigel (1986) site 516 (the 18 Ma Rio Grande Rise), and 597 (South Pacific) after Hart and Staudigel (1978, 1986) and the Troodos Ophiolite after Staudigel et al. (1986), Booij et al. (1995) and Gallahan and Duncan (1994). Vein mineral age cumulative curve is binned in 20 Myr age groups.

Oceanic-crust alteration involves formation of distinct compositional domains, which range in size from a ridge segment to a submilimeter sized vesicle filling. Obtaining a robust average of such chemically heterogeneous material is critical for understanding the chemical fluxes associated with seafloor alteration. Averaging heterogeneities with length scales smaller than the size of a typical geochemical sample is relatively easy, but becomes more difficult as the size of the heterogeneity increases. For large-scale heterogeneities, multiple samples need to be taken and characterized with respect to the proportion each sample contributes to the average. Such samples may be mixed in the correct proportions to produce composite samples that represent an average for a given section or crust type, or they may all be analyzed individually and then averaged arithmetically. Large-scale geochemical sampling requires a strategy that bridges the gap between sample size and the scales of key heterogeneities. Heterogeneities on the order of several meters to tens

Determining the Altered Composition of the Oceanic Crust of meters are probably the smallest relevant large-scale heterogeneities that need to be averaged. A compositional domain of this size is relevant geochemically because it may reflect a single chemical system, with its own characteristic phase assemblages that display consistent behavior during dehydration or melting events. This length scale is also a practical size for analyses of compositional domains through local observation of continuous exposure or core sections. Data from such a length scale can be relatively easily extrapolated to larger scales by visual integration of (.kilometer scale) cores, or field outcrop areas. The selection of individual samples for largescale compositional study may follow different strategies, depending on the type of crust studied and personal preferences. One of the major differences in strategy is the use of composite samples, versus the analysis of individual samples. Procedurally, these two approaches are identical, as long as the sampling strategy is comprehensive and the modal abundance data are determined in a consistent fashion. The two approaches each have advantages and disadvantages. Analysis of individual samples has the advantage that it offers the potential for resolving the chemical properties of different alteration types, before the samples are combined arithmetically into one average. The disadvantage of this approach is that it requires a large number of analyses and may be too labor intensive for some geochemical parameters. The advantage of composite samples is the significantly lower number of analyses required, which makes it much more likely to obtain a complete geochemical characterization of a representative set of samples at a particular crustal section. However, it is also possible to find a compromise between the use of composite or the individual samples. Staudigel et al. (1995, 1996) mixed one “Super” composite that represents the average of three drill sites, but they also analyzed two types of “subcomposites” to understand uncertainties and the internal structure of the “Super” composite. One type of subcomposite contrasts the compositions of volcaniclastics with flows that allow to contrast extreme with moderately altered regions. In addition, they made independent “depth” composites for two sites with similar alteration types to obtain an estimate of how variable such estimates can be for different sections. These subcomposites provide some insights into the chemical makeup of a site, without having to analyze all samples individually. Alternatively, it is possible to analyze all individual samples for geochemical parameters that are relatively easy to determine, and using composites only for the more involved analytical steps. The latter was done for all the highly altered

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volcaniclastic samples in the composites from sites 417 and 418 (Staudigel et al., 1996). Strategies for selection of individual samples may also vary. One approach is to use relatively large, representative samples that include veins, vugs, and variously altered haloes or other materials (Staudigel et al., 1989, 1995, 1996). Because of the complexity of alteration and the large number of permutations of primary and secondary features, several samples are typically required to approximate the chemical inventory of a particular compositional domain. This approach has the benefit of averaging out local heterogeneities or chemical gradients between veins and host rock, and minimizing uncertainties and contamination problems during sample handling. Another approach is to separate all major alteration types and vein materials, study them individually and recombine them in their respective proportions (e.g., Alt and Teagle, 1999; Bach et al., 2001). The advantage of this method is that it sheds light on the alteration processes and their impact on chemical fluxes more fully than composite samples do. The disadvantage is that it produces larger uncertainties in the overall averages, and requires substantially more sample handling and analyses. Any attempt to produce a bulk compositional estimate of altered ocean crust is critically dependent on accurate determination of the proportions of different rock types in the volume of interest. Simply averaging all analyses published for a particular drill site does not yield a realistic bulk compositional estimate for two reasons. First, low recovery during drilling biases the average towards the least altered material, whereas high recovery rates include a greater proportion of altered material. A hypothetical case of several holes drilled into the same crust at different recovery rates will show a positive correlation between degree of alteration and recovery rate. Secondly, the choice of samples for a particular study depends strongly on the scientific goals of that study. Sampling for igneous geochemistry studies typically focuses on the least altered samples. Such studies often further bias the sample selection by crushing the rocks and selecting the freshest rock chips out of a sample for analysis, and leaching carbonates out of the rock to obtain a maximally pristine chemical composition of an igneous rock. Such studies are obviously not useful in constraining the amount of alteration at a site. On the other hand, studies of alteration mineralogy of the crust focus their sampling on highly altered sections, especially extreme alteration types. Thus, a simple average of published geochemical data from a particular site is likely to offer insights mostly into the recovery rate and the types of studies performed rather than the actual average composition of oceanic crust at a particular site.

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Hydrothermal Alteration Processes in the Oceanic Crust

3.15.4 CHEMICAL CHANGES IN ALTERED CRUST COMPOSITION DUE TO HYDROTHERMAL PROCESSES 3.15.4.1 Time Dependence of Crust Hydration and Carbonate Addition Uptake of water and CO2 is one of the most sensitive indicators of alteration, and produces some of the most profound chemical changes in altered oceanic crust. These chemical changes are cumulative and offer the opportunity to evaluate the age dependency of alteration in drill sites of various ages. Presently, there are three sites in the upper oceanic crust that can be used for such comparisons: the 3.5 Myr old DSDP site 332B, the 5.5 Myr old site 504B, and the 120 Myr old sites 417A, 417D, and 418A (Figures 4(a) and (b)). Sites 417 and 418 are in oceanic crust that is substantially older than the predicted duration of hydrothermal alteration derived from heat-flow data (Figure 3) and thus these sites represent mature oceanic crust, which has experienced the complete history of hydrothermal alteration. Crust penetrated in site 504B is much younger than that in sites 417/418, and is in the age range of crust that is expected to have active hydrothermal alteration, based on the global heat-flow data set, even though local heat-flow measurements suggest that alteration is completed at this site

0

2

4

wt.% H2O 6

(Langseth et al., 1988). Crust of site 332B is even closer to the mid-ocean ridge and is expected to be experiencing active hydrothermal circulation. Unfortunately, sites 332B and 504B have very low recovery rates, and for this reason, some effort is needed to determine the true differences in time-variant alteration behavior. The highest contents of H2O are found in the upper 300 m of crust from sites 417A, 417D, and 418A (Figure 4(a)). At the same depth, the youngest crust at site 332B shows the least intense hydration, while crust at 504B has intermediate water content. Water contents of crust from 417/418 decrease with depth, while water contents at 332B and 504B increase slightly with depth before decreasing again. All trends converge at depths greater than 300 m. Differences in recovery rates is an unlikely cause for the trends observed in the upper 500 m of this diagram, and it is most likely that 332B and 504B have not reached a mature degree of hydration in their upper portion, while 504B is likely to have completed its hydration at depths . 600 m. Thus, the formation of layer silicates in the upper portion of 504B and 332B is probably an ongoing process. CO2 shows behavior similar to water. Its concentration is also highest in the upper 300 m of crust at sites 417 and 418 and lowest in 504B throughout the depth intervals covered by all holes. It is interesting to note that crust of site wt.% CO2

8

2

0

6

4

0

Depth (m)

200

400

600 504B 332B 417 / 418

800

Super 417D 418A

1,000 (a)

(b)

Figure 4 (a) Water versus depth and (b) CO2 versus depth for major deep drill sites in the Atlantic and Pacific Oceans. Data compilations are taken from the Initial Report Volumes of the Deep Sea Drilling Project or the Ocean Drilling Project (site 332B: Aumento and Melson, 1977; 417/418 Donnelly et al., 1979b; site 504B: Alt et al., 1993; bold lines connect depth composite data from sites 417D and 418A (after Staudigel et al., 1995).

Chemical Changes in Altered Crust Composition Due to Hydrothermal Processes 332B has higher CO2 values than that of 504B, particularly in the depth interval 400– 500 m, even though it samples younger crust. With increasing depth, CO2 values of 417/418 crust show a systematic decrease to values that are substantially above the values from site 504B for some depth. These observations indicate that CO2 uptake in oceanic crust is highly variable in younger crust, and both young crustal sections are distinct from mature ocean. Site 504B appears to have particularly low CO2 abundances. This, combined with incomplete hydration, suggests that 504B crust has not experienced all of the carbon uptake it will see throughout its lifetime, and for that reason it is not a reliable reference point for extrapolating to global mass balances. This is acknowledged in the global carbon budget of Alt and Teagle (1999) who ignored the data from the upper portion of 504B. 3.15.4.2 Chemical Fluxes between Oceanic Crust and Seawater: Methods and Uncertainties Geochemical exchange between seawater and the oceanic crust varies substantially as a function of depth and lithology, and for this reason alteration processes are considered separately for different portions of the crust in the 7.1 km oceancrust reference section in Figure 2. Chemical fluxes are given as D values, the difference in measured and original chemical inventory in g/100 g fluxes (at constant Ti), independently for each depth interval in the oceanic crust. An average altered MORB may be obtained by adding this flux to the fresh MORB listed in Table 2. The flux estimates in Table 2 primarily derive from the data from sites 417D and 418A for the upper crust, and site 735B for gabbros of the deep crust, because these are the only drill sites with high recovery rates and total depths greater than 300 m that sample crust older than 10 Ma, which is considered a minimum age for “mature” or near-mature ocean crust alteration. Study of sites with a critical depth of penetration (.100 m?) is more likely to help us understand their true behavior than studies of very shallow sites where it is difficult to establish the overall alteration behavior. In particular, sites 417A, 417D, and 418A allow for an in-depth understanding of localized anomalies. For sites 417/418, the fluxes are calculated from the “super” composite, which is dominantly composed of materials from sites 417 and 418A and offer a grand average for these sites. In its upper 200 m, this composite includes about 20% of the oxidatively weathered materials from site 417A. Sites 417/418 and 735 are the only deep drill holes displaying average recovery rates .70%, allowing a quantitative reconstruction of

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the alteration inventory. The main sources of data for these sites come from Staudigel et al., 1989, 1995, 1996; Smith et al., 1995; Spivack and Staudigel, 1994; Hart et al., 1999; Bach et al., 2001. However, for the intermediate crustal sections, data from site 504B was also utilized (e.g., Alt et al., 1993a,b; Alt and Teagle, 1999; Chan et al., 2002). This procedure was adopted because there are no other data from in situ crust. It is important to note that use of these data leads to increased uncertainties as alterations may not be complete and the recovery rates are very low for 504B. The results from these deep-sea cores are also compared to evidence from the Troodos and Samail ophiolites (Bednarz and Schmincke, 1989, 1994; Bednarz, 1989; Bickle and Teagle, 1992; Spooner et al., 1977). However, in many cases, fluxes were extrapolated between site 735 and Sites 417/418, using ophiolite data or data from 504B only for comparison. This procedure was necessary for a variety of reasons including lack of data, incomplete alteration due to young oceanic-crustal age, and to produce the most conservative estimate. Obtaining realistic errors is one of the most difficult, yet most crucial problems in all flux estimates. Such errors can be approximated through an independent error analysis for several factors that are involved in estimating fresh and altered rock composition. There are uncertainties arising from petrographic observations, in the choices of representative samples, recovery rate biases, and analytical errors. In most cases analytical errors are a relatively minor source of uncertainty, and they are typically rather well documented. Probably the most crucial analytical uncertainty is in acurately determining the titanium concentration that is used as a normalizing factor to account for open-system behavior. This uncertainty directly relates to an error in the fluxes, and thus fluxes are difficult to constrain to better than ,1% of the whole rock abundance of a particular element. Uncertainties in phenocryst abundance estimates are a significant source of error, particularly for major and highly compatible elements in the main phenocryst phases. In most cases related to ocean drilling samples, phenocryst abundances in a particular sample are rarely known to better than about 5 vol.%. The most abundant phenocrysts are olivine and plagioclase, which contain the elements magnesium, silicon, calcium, aluminum, and nickel. Detailed uncertainty estimates for major elements are given in Staudigel et al. (1996, table 6). An indication of the overall reproducibility of flux estimates can be obtained by comparing independent data from different sites having similar alteration behavior, like site 417D and 418A. The H2O and CO2 contents for

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Hydrothermal Alteration Processes in the Oceanic Crust

composites at sites 417D and 418A as a function of depth are compared to the large scale, bulk composition at these sites in Figure 4. CO2 abundances at these two sites track each other very well, with an uncertainty between these sites of ,0.5%. In contrast, the two H2O curves vary more widely (^1%). Close agreement between composites from these sites suggest a high degree of reproducibility, while a large scatter suggests substantial variation between sites or a problem with reproducibility. In the following analysis this scatter within and between sites is used to indicate how robust these flux estimates are for the upper oceanic crust.

3.15.4.3 Chemical Fluxes H2O. When discussing the water contents of the oceanic crust it is important to know that it may be found in three forms that are measured differently and that are relevant for different applications. Ocean crust contains (i) formation waters filling pore spaces, (ii) exchangeable water in minerals like clays or zeolites, and (iii) water bound to crystalline structures or dissolved in silicate glass. Formation water can be estimated from the pore space available in the formation, which is typically done by borehole logging. Exchangeable water (H2O2) is typically defined as the water expelled at 110 8C, even though the release of this type of water is rather continuous and some fraction may occur at temperatures higher than this. Crystal-bound water is referred to as H2Oþ and it is determined in rock powders by titration or chromatography following ignition at temperatures .800 8C, or, in situ by spectroscopic methods. In most geochemical studies H2O is understood to be H2Oþ, in particular when H2O2 is not reported. Non-degassed, fresh MORB has approximately 0.2 wt.% H2O. The amount of formation water in the upper crust probably scales directly with the porosity of the crust, as it can be determined by borehole logging. The upper crust has about 10 vol.% porosity (e.g., 13% in site 417D; Salisbury et al., 1979), which would yield roughly 4 – 5 wt.% formation water for the upper crust. There are insufficient numbers of H2O2 measurements to estimate the loosely bound water in these sections. Staudigel et al. (1995) estimated a H2Oþ flux for the upper crust of about 2.81 g/100 g. It is interesting that the average H2Oþ of composites from the very top of cores 417D and 418A have values that are significantly below the values at 100– 200 m depth, and that site 418A has values that are substantially above the values of 417D, giving an approximate uncertainty of ,0.5 wt.% for the supercomposite. Taking this average H2Oþ

as typical of the upper crust and including the estimate of formation water given above yields a total of 7 –8 wt.% H2O. Assuming ,2 –3 wt.% for H2O2 suggests that the upper oceanic crust stores close to 10 wt.% water, a water content that is not much different from those of pelagic sediments. This value is likely to decrease substantially with depth, because pore space decreases with depth and because higher-temperature metamorphic assemblages contain less water. An estimate of water content of the gabbro section can be made using data from site 735B. Bach et al. (2001) suggest that gabbros from site 735B took up about 0.11 g H2Oþ/100 g of rock. Formation water can be constrained from porosities in this site. Core materials from site 735B have porosities of 0.7 vol.%, and in situ porosity is likely to be twice as high, probably about 1.5 vol.% (Dick et al., 2000), which is substantially less than that of the upper crust. Filling 1.5 vol.% of the rock with water yields about 0.5 wt.% of H2O, suggesting that the gabbros are likely to store slightly less than 1 wt. % water—an order of magnitude less than that present in the upper crust. The mass balance uses the H2Oþ flux of Bach et al. (2001), which should be close to a total average for the lower crust (even though 735B penetrated only through its upper portion). Water contents at intermediate depths are extrapolated from the top of the crust downward using the internal variation in sites 417/418 and the data for minimal alteration from site 504B and site 735B. The H2Oþ fluxes into the lower extrusives and sheeted dikes are estimated at 2.10 g/100 g and 1.5 g H2Oþ/100 g, respectively (see Figure 4(a)). Adding up and averaging these individual estimates, the oceanic crust as a whole takes up ,0.45 g H2Oþ/100 g of rock. The total water inventory of the whole crust (including H2Oþ, H2O2, and formation water) is derived simply by scaling (in a similar fashion as in the upper crust), at ,1.5 wt.%. CO2. Degassed MORB has ,0.12 wt.% CO2, whereas undegassed MORB has ,0.45 wt.% CO2 (Table 2; Gerlach, 1989). The higher CO2 in altered MORB is due to CO2 addition to the crust from seafloor alteration subsequent to initial outgassing. This alteration-related addition of CO2 to the oceanic crust brings its total inventory up to levels exceeding the original non-outgassed carbon contents (Staudigel et al., 1989). Most of this hydrothermal CO2 inventory in the oceanic crust is added in the form of carbonates, particularly in the upper 600 m (Figure 3(b), Staudigel et al., 1989). In the estimate of carbon fluxes (Table 2), data for the upper 600 m of sites 417/418 yielded a CO2 uptake of 3.26 g/100 g (Staudigel et al., 1989). Based on the scatter of the depth composites and the discrepancies between sites 417D and 418A, the error on these

Chemical Changes in Altered Crust Composition Due to Hydrothermal Processes estimates is roughly 0.3 g/100 g. This estimate from sites 417 and 418 was confirmed by Alt and Teagle (1999) using data from sites 843 (90 m penetration) and 801(150 m penetration), which yielded 2.4 wt.% and 3.55 wt.% CO2., respectively. For the lower crust we use carbon uptake estimated by Bach et al. (2001) for the gabbros at site 735B, yielding a CO2 flux of 0.06 g/100 g into the 1,500 m gabbro section. Lower extrusives and sheeted dikes from site 504B apparently have not seen the complete cycle of alteration, and therefore the very low CO2 values from these sites are not robust indicators of the total carbon budget of the intermediate crust. To estimate the carbon inventory of the intermediate levels of the oceanic crust, we interpolate between the extrusives and the gabbros, which yields an estimate of 0.5 g CO2/100 g for the lower extrusive section, slightly above the value of Alt and Teagle’s (1999) transition zone of site 504B. For the sheeted dikes we use 0.06 g/100 g, the value of the gabbro section, as a conservative minimum estimate. The total uptake of the crust is calculated here as 0.355 g/100 g, which is slightly above the estimate made by Staudigel et al. (1989), and also similar to the total carbon uptake of the oceanic crust inferred by Alt and Teagle (1999) and for the Troodos ophiolite (Bednarz and Schmincke, 1989). The above extrapolation of data for intermediate depths (lower extrusive crust, sheeted dikes) does not contribute major uncertainties to the total flux estimate, because most of the carbon inventory is located in the upper 600 m of the crust, and, therefore, most of the uncertainties lie in this depth interval. Oxygen isotopes. d18O in ocean-crust studies is typically defined as the per mil deviation in 18O/16O ratio of a rock relative to a standard mean ocean water (d18OSMOW) and it is widely used to understand ocean-crust alteration processes. Fresh MORB has an d18OSMOW value of þ 5.7‰, and water – rock interaction with seawater (d18OSMOW ¼ 0‰) at low temperatures increases the value, while high-temperature alteration decreases it. Muehlenbachs and Clayton (1972) drew attention to this relationship and suggested that hydrothermal alteration of the crust may buffer the oxygen isotopic composition of seawater. Oxygen is the major component in the oceanic crust, and therefore, changes in d18O are a rather profound indicator of hydrothermal alteration. The silicate portion of the upper oceanic crust at sites 417/418 has a d18O of ,þ10‰ (þ 9.98‰, Staudigel et al., 1995), substantially elevated from the original magmatic value of þ 5.7‰. The values for the composite crust decrease with depth from ,þ11.5‰ in the upper 100 m, down to ,þ 8.5‰ between 400 m and

525

600 m depth, roughly defining an uncertainty band of ,^1d18O units. This uncertainty is rather small when compared to data from other oceancrust sections. For example, the same depth interval in the Troodos and Samail ophiolites displays values ranging from ,8‰ to 14‰ (e.g., Spooner et al., 1974; Gregory and Taylor, 1981; Stakes and Taylor, 1992), which is substantially larger than the 417/418 composite range, but these ophiolites also display a general decrease of d18O with depth. Site 735B shows a rather interesting variation, even though the data are more scattered. Its uppermost 500 m display a range of 3– 7‰, with an average d18O of 4.35‰ (Hart et al., 1999), well below the value of fresh MORB, and clearly indicating relatively high reaction temperatures. Deeper sections of 735B display heavier d18O values, also with a significant scatter, but with a higher average value, on the order of 5.8‰. This is not very different from the original magmatic value (Bach et al., 2001). This pattern of very low d18O values at the top of the gabbros that slightly increase with depth is quite similar to the variation seen in the Troodos and Samail ophiolites (Spooner et al., 1974; Gregory and Taylor, 1981; Stakes and Taylor, 1992). The d18O of the average crustal section in Table 3 derives from the average d18O for the top 600 m of sites 417/418 (d18O ¼ 10‰), with d18O decreasing to 8‰ in the lower extrusives, 4‰ in the sheeted dikes and then increasing again to 5.8‰ in the gabbros from site 735B. This overall variation is solely derived from drill holes in normal oceanic crust, but it is remarkably similar to the patterns observed in ophiolites. This suggests that 735B is actually quite an appropriate analogue for hydrothermal alteration in “true,” deeply buried gabbro sections in normal oceanic crust. It appears that most hydrothermal systems penetrate into the top few hundred meters of gabbros, whether or not the section carries any overlying extrusives and dikes. Strontium. The 87Sr/86Sr ratio is also a very powerful indicator of hydrothermal alteration in the oceanic crust, because fresh crust is unradiogenic (87Sr/86Sr ¼ 0.7025 –0.703; Chapter 3.13) while seawater is radiogenic (0.7092; Chapter 6.02). Seawater 87Sr/86Sr also changes with time. One-hundred-million year old seawater was less radiogenic (0.70735) than today’s, and 87Sr/86Sr has been continuously increasing since then. The isotopic composition of reaction products quantitatively reflects the strontium contributions from seawater and from basalt. The reference sites 417/ 418 and 735B both display primary strontium ratios close to 0.7029 (Staudigel et al., 1981; Hart et al., 1999; Bach et al., 2001). The average 87 Sr/86Sr ratio at sites 417/418 is 0.704575, with a rather systematic down-hole variation that closely

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Hydrothermal Alteration Processes in the Oceanic Crust

mirrors the behavior of H2O. Alteration-related increases in 87Sr/86Sr at the top of the crust are relatively modest, become more pronounced at a depth of ,200 m, and then decrease with depth again, to its lowest values at 550 m (Staudigel et al., 1995). Hart et al. (1999) and Bach et al. (2001) showed that site 735B has generally very low 87 Sr/86Sr (0.70295), with some small-scale variation that might resemble several cycles of decreasing 87Sr/86Sr with depth. 87Sr/86Sr values from sites 417/418 are high relative to those from site 504B (e.g., Barret and Friedrichsen, 1982; Friedrichsen, 1984; Kawahata et al., 1987), but are relatively low when compared with 87Sr/86Sr of the Troodos ophiolite (Bickle and Teagle, 1992). The former might be due to incomplete alteration at 504B and the latter due to higher initial isotopic ratios in fresh basalts in the Troodos ophiolite. The trends in 87 Sr/86Sr at sites 417/418 and in the Troodos ophiolite are used to extrapolate the 87Sr/86Sr of the lower extrusives (87Sr/86Sr ¼ 0.704) and the sheeted dikes (0.7035). For the gabbros, 87Sr/86Sr is assumed to be 0.70295, which is only slightly elevated from the original values (Hart et al., 1999; Bach et al., 2001). Strontium isotope ratios and abundances in samples from the oceanic crust may be used to determine the complete chemical mass balance of strontium exchange between seawater and basalt, including the loss of basaltic strontium to hydrothermal solutions, and uptake of basaltic or seawater strontium from hydrothermal solutions. A mass balance of this exchange can be made in four steps. (i) The relative amount of basaltic and seawater strontium in altered basalt can be determined from the measured 87Sr/86Sr of an altered sample as a (linear) mixture of strontium from the two end members, the (contemporaneous) seawater and basalt. (ii) The inventory of the basaltic and seawater strontium in an altered sample (in mg/kg) may then be determined from the above ratio of seawater and basalt in the sample and the total strontium abundance measured for this sample. (iii) Seawater strontium addition to the basalt is given directly by the seawater strontium inventory calculated in Step (ii). (iv) The determination of flux of basaltic strontium in or out of an altered sample is more complicated because it has to be related to the original inventory of strontium. It is determined as the difference between the original basaltic inventory and the basaltic strontium present in the altered sample. Using this type of mass balance, it can be shown that on average, basalts from the upper crust at sites 417 and 418 have lost 10 mg/kg of basalt strontium to seawater and gained 32 mg/kg of seawater strontium, resulting in a total flux of 22 mg/kg of strontium into the basalt (Staudigel et al., 1996). Gabbros at 735B show very large

changes in strontium abundance between fresh and altered sample pairs, but these changes show both gains and losses of strontium by up to 25% of the total inventory (Bach et al., 2001). Hart et al. (1999) estimated from strip composite samples that the upper portion of 735B shows no or very little change in strontium abundances. In Table 2, strontium fluxes are derived from the data of sites 417/418 and 735B for the upper and lower crust, respectively, effectively tapering the rather high strontium fluxes into the upper crust to the near zero strontium fluxes in the gabbros. A slight strontium loss is assumed for the gabbros, because the breakdown of calcic plagioclase is likely to liberate substantial quantities of strontium, some of which may be removed by hydrothermal solutions. Over the total crust, these elemental changes add up to a slight uptake of strontium, of 1.4 mg/kg of rock, which is extremely small when compared to the overall fluxes of strontium into and out of the crust. Thus, seafloor alteration can substantially influence the isotopic composition of seawater, but not significantly change the strontium inventories in either reservoir (see Palmer and Edmond, 1989; Elderfield et al., 1999). Sulfur. Sulfur is a key element in hydrothermal processes. Seawater introduces sulfate into the hydrothermal system that precipitates anhydrite during heating (e.g., Shanks et al., 1981). Basalt is very reduced and contains sulfur dissolved in melt and immiscible sulfur metal globules (Mathez, 1976, 1980). The reaction of seawater with basalt results in the precipitation of sulfides, mostly pyrite, as one of the most common reaction products. There are only very few systematic sulfur and isotope studies of the oceanic crust, in particular those by Alt (1995) on ODP site 504B, Bach et al. (2001) on site 735B, and Alt (1994) on the Troodos ophiolite. The pre-alteration sulfur abundance at site 504B is 960 ppm. Overall, basalts appear to be a source of sulfur to hydrothermal solutions, delivering ,0.1 –2 g of sulfur per 100 g of basalt (recalculated from Alt, 1995) and sulfur is precipitated in veins (sulfide and anhydrite), in particular in mineralized zones at the bottom of the extrusives and top of the dikes (Alt, 1995). Alt (1995) determined the volumes of veins in 504B and estimated a total flux of 0.15 £ 1012 g yr21 from the oceanic crust, cautioning that this is the only site where such a mass balance has been done and citing relatively large errors. However, the total sulfur flux in altered oceanic crust is relatively small when compared with the sulfur in high-temperature hydrothermal fluxes ((10.3– 25.7) £ 1012 g yr21; Von Damm et al., 1985; Sleep, 1991) and river fluxes (28.5 £ 1012 g yr21; Holser et al., 1988). Alt (1995) suggests that the discrepancy in the hydrothermal fluxes is probably best explained

Chemical Changes in Altered Crust Composition Due to Hydrothermal Processes by anhydrite precipitation during black-smoker activity, and subsequent dissolution from lowtemperature fluids. In Table 2 a sulfur loss of 0.06 g/100 g (Alt, 1995) is estimated for the extrusive section and a sulfur gain of 0.02 is estimated for the gabbros at site 735B (Bach et al., (2001). The total sulfur flux is 0.0037 g/100 g from basalt to seawater. The d34S of fresh MORB is þ 0.1‰ (Sakai et al., 1984). The calculated d34S of 21.9‰ for the extrusives, þ1.3‰ for the dikes (anhydrite veins not included) and þ 0.13‰ for the gabbros (Table 2) is taken from Alt (1995), who estimates the average d34S for the total oceanic crust as þ 0.9‰ per mil. Lithium. Lithium is a very sensitive indicator of hydrothermal alteration processes, because it is added during low-temperature alteration, mostly in smectites, and is leached during hightemperature alteration (e.g., Seyfried et al., 1984; Von Damm et al., 1985). In addition, lithium isotopes reflect seawater additions to the upper crust, much like strontium and oxygen isotopes do. The most comprehensive study of lithium abundances and isotopic ratios was recently presented by Chan et al. (2002) for site 504B. Considering the low recovery rates at this site, this study should be considered to provide a minimum baseline of lithium isotopic alteration, with an emphasis on the early portion of alteration. Unaltered MORB has an average lithium abundance of 4.5 ^ 1.5 ppm (Chan et al., 1992; Ryan and Langmuir, 1987). The least altered rock from site 504B has ,3 ppm lithium, the average lithium inventory in the volcanics is ,5.8 ppm, indicating an uptake of 2.8 mg/kg. The average lithium content of the sheeted dike (þ the transition zone) is ,2.5 ppm, indicating a loss of 0.5 mg/kg. Chan et al. (2002) suggest that the initial lithium content of site 735 is ,3 ppm and the measured value is 1.0, suggesting a loss of 2 mg Li/kg. In total, the crust loses ,1.09 mg Li/kg of basalt. d 7Li, the per mil deviation of the 7Li/6Li ratio from the NBS SRM L-SVEC, of fresh MORB is þ3.4 – 4.7‰ (Chan et al., 1992; Ryan and Langmuir, 1987) and rises steeply with the degree of low-temperature alteration (Chan et al., 1992), reflecting uptake of heavy seawater lithium (þ32‰). d 7Li decreases with depth in site 504B, from an average of ,6‰ in the extrusives to ,2‰ in the sheeted dikes. This reflects addition of seawater lithium to the upper crust and removal and fractionation of lithium in the deeper crust (Chan et al., 2002). Boron. Boron concentrations in fresh MORB are ,0.5 ppm (Spivack and Edmond, 1987; Ryan and Langmuir, 1987) and 26.2 ppm in average upper crust at sites 417 and 418 (Smith et al., 1995). Boron concentrations in altered crust show

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a significant decrease with depth, from ,50 ppm in the upper 100 m of 417/418 to 8.5 ppm at 500 m (Smith et al., 1995). Subtracting the initial basaltic boron inventory from the altered extrusive inventory yields a boron uptake of 25.7 mg/kg for the upper 600 m of extrusives. Extrapolating the 417/418 data downward, the lower extrusives have ,4.5 ppm boron, and an uptake of 4 mg/kg (Table 2). An uptake of 5.6 mg/kg for the sheeted dikes and 2.2 mg/kg for the gabbros is estimated using Layer 2B and Layer 3 estimates for boron from Smith et al. (1995; 5.9 ppm and 2.7 ppm, respectively). The total uptake is thus 4.81 mg/kg for the whole oceanic crust. d11B (i.e., the per mil deviation of the 11B/10B ratio from the NBS SRM 951 standard) in the upper crust can increase substantially during seafloor alteration. Hyaloclastites may reach d11B values up to þ5.4‰, but the average crust is only slightly elevated (þ 0.8‰) relative to the least altered composites of flows (0.5‰, Smith et al., 1995). Gabbros from the upper portion of site 735B have markedly elevated d11B of 7.35‰ (Hart et al., 1999). d11B for intermediate depths are extrapolated from these values to be 1‰ for the lower extrusives and 2‰ for the sheeted dikes, resulting in an overall value of d11B ¼ 5.6‰ for the total crustal section. Potassium, rubidium, and caesium. They are amongst the most sensitive indicators of water– rock exchange. Like their sister alkali element, lithium, these elements are taken up by basalts during low-temperature alteration reactions and leached from basalts during high-temperature reactions. The upper 600 m of crust at sites 417 and 418 records gains of 0.54 g/100 g K2O, 10.3 mg/kg rubidium and 0.183 mg/kg caesium (Staudigel et al., 1995, 1996; Hart and Staudigel, 1989). Surprisingly, the gabbros at site 735B all show gains of these alkalis as well: 0.01 g/100 g K2O, 0.7 mg/kg rubidium and 003 mg/kg caesium (Bach et al., 2001). This may reflect the anomalous nature of the crust at 735B, because it is directly exposed to seawater, or it may be a “typical” latestage low-temperature alteration overprint that affects all of the oceanic crust, even if it was previously depleted in potassium, rubidium, and caesium from high-temperature reactions at the ridge axis. For this reason it is not clear whether the lower oceanic crust displays slight gains or nearcomplete losses in these elements. Given the rather low initial inventories of potassium, rubidium, and caesium in oceanic gabbros and the very large gains in the upper crust, the above uncertainties in the gabbro inventories have a relatively small effect on the alteration budget of these elements for the whole crust. For these reasons, the data from sites 417/418 are used for the upper oceanic crust and depth trends at these sites are used to extrapolate the values for the lower 400 m of the extrusives. The sheeted dikes and the gabbros are

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Hydrothermal Alteration Processes in the Oceanic Crust

proposed to have experienced no net change in potassium, rubidium, and caesium. Using average values of 0.05 g/100 g K2O, 1 g/kg rubidium, and 0.03 g/kg for caesium results in total fluxes of 0.0485 g/100 g of K2O uptake, 0.927 mg/kg of rubidium uptake, and 0.0172 mg/kg of caesium uptake for the whole crust. Uranium. Uranium content is also enhanced in the upper oceanic crust during alteration. The uranium content of the altered upper crust is 0.30 ppm (Hart and Staudigel, 1989; Staudigel et al., 1996). Uranium is slightly depleted in the upper part of site 735B (Hart et al., 1999), but slightly enriched overall in 735B (Bach et al., 2001). Data for the top 600 m of site 417/418 are therefore used for the upper crust, and the grand average of 735B is used for the lower crust. As in the previous estimates, the internal variation in 417/418 and 735B are used to extrapolate the intermediate crust, yielding a total uptake of 0.037 mg/kg of uranium in the crust. Rhenium and osmium. There are no data on behavior of Re/Os during alteration of MORB, but some work has been done on pillow lavas from the submarine seamount sequence in the basement complex of La Palma (Canary Islands; Marcantonio et al., 1995). There, altered pillow rinds have 187Os/188Os ratios of 0.40– 0.43, while the pillow interiors generally have 187Os/188Os values between 0.14 –0.19. Schiano et al. (1997) found highly variable 187Os/188Os ratios in midocean ridge glasses and cannot rule out potential assimilation of manganese crusts with high 187 Os/188Os. Gabbros from site 735B suggest isotopic variations consistent with these changes (Hart et al., 1999). Thus, seafloor alteration is likely to increase the 187Os/188Os significantly in some rocks, in particular in the upper crust (Marcantonio et al., 1995), but it is too early to determine whether this effect is of any more than localized importance. SiO2 is amongst the most difficult element to track during alteration of the oceanic crust because it shows major variations due to igneous processes and because the majority of alteration reactions involve silicates. The upper extrusive crust at 417/418 gains about 1.2 g/100 g SiO2. The uncertainties are expectedly very large, including ^0.18 g/100 g in the estimate of the fresh rock composition, and a variance of nearly ^4 g/100 g in fluxes at various depths of sites 417D and 418A (Staudigel et al., 1995, 1996). Bach et al. (2001) estimated that gabbros of site 735B gained about 0.1 g SiO2/100 g gabbro, and the scatter in their data suggest that this value is also rather uncertain. We extrapolate between these values and estimate the fluxes to be 0.5 g/100 g for the lower extrusive crust and 0.25 g/100 g for the dikes (Table 2). However, these values are likely to have errors on the order of ^2 g/100 g, which are extremely large

when compared to fluxes at hydrothermal vents and in the ocean in general. Al2O3. Fresh MORB contains, on average, 15.26 wt.% Al2O3 (Hofmann, 1988; see also Chapter 3.13). Alteration of the upper crust causes a gain of about 1.94 g/100 g of Al2O3 (Staudigel et al., 1996), based on individual fluxes at 417D and 418A that vary from 21 g/100 g to 5.0 g/100 g. This range indicates with significant uncertainty, partly due to the large uncertainty in the fresh estimate (^0.37 wt.%, mostly from uncertainties in plagioclase phenocryst distribution) and to local redistribution of aluminum within the holes. An analytical uncertainty of ,1% in the estimate of the altered rock composition would contribute an error of ^0.15 g/100 g to the fluxes. The lower crust Al2O3 content is unlikely to be affected by alteration. FeOtot in unaltered crust is 10.43 wt.% (Hofmann, 1988; see also Chapter 3.13). The upper crust at 417 and 418 loses about 0.01 g/ 100 g; individual composites may show losses to 0.85 g/100 g or gains of up to 1.01 g/100 g of up (Staudigel et al., 1996). These variations are most likely to represent uncertainties in the estimates rather than true fluxes. Iron concentrations in MORB are very high and variable while seawater contains very little iron. The analytical uncertainty in the determination of iron alone is on the order of 0.1 wt.%, which is almost an order of magnitude higher than the flux. This shows that iron is essentially immobile and that the fluxes are too low to detect beyond the uncertainties of crustal mass balances. However, due to the extremely lowiron inventories in seawater, even the smallest fluxes would be quite significant for the seawater budget, and, thus, the oceanic crust may deliver substantial fluxes to seawater. Given the concentration gradient, these fluxes are likely to move iron from the crust to the oceans but there is no convincing data from the oceanic crust that can prove this. MnO shows no change in the upper extrusives, but fluxes for individual composites range from 0.04 g/100 g to 0.05 g/100 g (Staudigel et al., 1996) suggesting very large uncertainties that do not allow us to estimate meaningful fluxes independently. The very low manganese concentration in seawater and the enrichment of manganese in hydrothermal waters suggests that it is likely that manganese leaves the crust, even though local enrichments of manganese are likely near the top of the extrusive oceanic crust due to deposition of manganese crusts and Fe – Mn hydrothermal deposits. MgO is a key element in the discussion of hydrothermal systems. It is an important element in many igneous phases and in alteration phases. The depletion of magnesium in hydrothermal vent fluids was originally seen as evidence that the

Discussion oceanic crust is a major sink for magnesium from riverine input to the oceans (Edmond et al., 1979; Van Damm et al., 1985). However, flux estimates from rock analyses do not to support this contention, at least not in a very consistent manner. Data from sites 417/418 and 735B suggest that, on average, rocks lose more magnesium than they take up (0.2 g/100 g in both cases, Staudigel et al., 1996; Bach et al., 2001). Dredged samples of abyssal peridotite suggest that they lose up to 5 wt.% magnesium due to alteration (Snow and Dick, 1995). Magnesium uptake was reported at the Troodos ophiolite (Bednarz, 1989; Bednarz and Schmincke, 1989), and for site 504B (Alt et al., 1996). Part of this discrepancy may be due to the very large uncertainties in these fluxes. At sites 417/418, fluxes of individual depth composites vary from 21.1 g/100 g to þ2.01 g/100 g, and the uncertainty in the fresh rock estimate is 0.19 g/ 100 g. Thus, the data from 417/418 actually permit a small gain in MgO, but the uncertainties are large. The estimate in Table 2 uses the 417/418 data as representative of the upper oceanic crust and makes the most conservative estimate of lower crust/sheeted dikes of a magnesium flux of zero (Table 2). The crustal budget of magnesium is similarly as poorly constrained as the budget for iron, but it is likewise clear that there really are no very large fluxes. A major difference between magnesium and iron is that magnesium has a much larger inventory in seawater, which requires much larger fluxes to make it significant for the seawater – ocean-crust budget. These large fluxes are not observed. CaO. Calcium is also an element that is very important to global geochemical cycles, as the oceanic crust is considered to be an important source of Ca2þ (e.g., Berner et al., 1983). There are several competing processes that control the calcium content in the upper crust. Calcium is lost during glass alteration (Staudigel and Hart, 1983) and calcic feldspar breakdown, and is taken up during calcium carbonate precipitation. Sites 417/418 display an uptake of 2.1 g/100 g ^ 1 g/ 100 g CaO (Staudigel et al., 1996), with an overall decrease with depth. Alt et al. (1996) report a small loss of CaO at site 504B, which is consistent with the limited precipitation of carbonate there, while much of the calcic feldspar is replaced. Extrusives in the Troodos ophiolite drill cores consistently show loss of CaO: 8.08 g/100 g in the upper crust and 3.34 g/100 g in the lower crust (Bednarz, 1989). The flux values for CaO in Table 2 use site 417/418 data for the top of the crust, and site 735B data for the gabbros. These sites are considered the most representative of the oceanic crust. Intermediate values in the lower extrusives and sheeted dikes are extrapolated from these data and tapered to zero fluxes, like CO2,

529

largely because much of the calcium deposition appears to be linked to carbonate precipitation. This yields an uptake of 0.03 g/100 g in the lower extrusives and an uptake of 0.04 g/100 g in the sheeted dikes. Altogether these calcualtions suggest a slight positive flux of calcium into the crust of 0.134 g/100 g. Most evidence suggests that calcium contents do not show much net change, even though it is quite likely that the exchange of calcium between seawater and basalt is very high, probably quite similar to the behavior of strontium. Na2O is a mobile element that is involved in important alteration reactions in the shallow crust (loss during glass alteration, gain from formation of zeolite), in the deeper crust (formation of zeolites, albite, and halite) and has a substantial inventory in seawater, which is at times enhanced by boiling of water in high-temperature hydrothermal systems followed by the precipitation of halite. The oceanic crust appears to gain small quantities of sodium, on average, particularly in the upper crust (0.15 g/100 g in 417/418; Staudigel et al., 1996). Bach et al. (2001) suggest a gain of 0.03 g/ 100 g for the gabbros at site 735B. A gain of 0.15 g/ 100 g for the lower extrusives and the dikes is used in Table 2. This is intermediate between the smaller uptake at site 504B (0.09 g/100 g; Alt et al., 1996) and the higher uptake at the Troodos ophiolite (1.9 g/100 g lower extrusives; 0.421 g/100 g sheeted dikes; Bednarz, 1989). Like other major elements, however, it is obvious that the gains or losses of sodium in basalts are quite variable and uncertainties are high. A large number of elements are immobile during alteration of the oceanic crust. These are generally elements that are insoluble in seawater and include titanium, which is commonly used as an immobile reference element to constrain overall losses or additions in a chemically open system and REEs, hafnium, niobium, zirconium, and thorium. Similarly, the neodymium and hafnium isotopic composition of oceanic crust is rarely affected by alteration. However, significant mobility of otherwise immobile elements has been identified in extremely altered basalts dredged from some seamounts (e.g., Cheng et al., 1987). Staudigel et al. (1996) observed distinct but minor changes in 143Nd/144Nd and in Ce/Cep in the most altered composite samples from sites 417 and 418, consistent with the addition of REE in very large quantities of seawater or moderate amounts of sediment particulates (Staudigel et al., 1996).

3.15.5 DISCUSSION 3.15.5.1 Hydrothermal Fluxes: Rock Data versus Fluid Data The estimates of elemental changes due to hydrothermal alteration of oceanic crust in Table 2

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Hydrothermal Alteration Processes in the Oceanic Crust

can be re-cast into global geochemical fluxes using the ocean-crust production rate discussed above (see Figure 2). These fluxes can then be compared with river fluxes and two types of hydrothermal fluxes: axial hydrothermal fluxes or flank fluxes following Wheat and Mottl (2000). 3.15.5.1.1

Uncertainties

Errors in elemental fluxes derived from crustal estimates are larger than, or similar to, the value of the actual flux estimate for many of the major elements: silicon, aluminum, iron, manganese, magnesium, and sodium. The fluxes for these elements are thus poorly constrained, but these estimates do serve as conservative bounds on the fluxes. Unfortunately, these bounds overlap the fluxes derived from hydrothermal fluid data and river data. For this reason, current ocean-crust flux estimates do not provide independent evidence for the magnitude of hydrothermal fluxes in the geochemical cycle for these elements. Within the bounds of these uncertainties, the data indicate that ocean floor hydrothermal processes may balance (or compound!) missing global fluxes of these elements. For these reasons, these elements are not discussed here in any detail. The flux estimates of CaO are slightly better known than the forgoing major elements, but the uncertainty of CaO flux in the upper oceanic crust is still almost two-third of its value. Much larger uncertainties are typical of the deeper extrusives and sheeted dikes, where spilitization is characterized by near-complete loss of calcium. All other element fluxes reported here have uncertainties that are less than 30% of their value. 3.15.5.1.2

Bulk fluxes

A comparison can be made between fluxes derived from the rock record versus those obtained from hydrothermal fluids for elements for which well-constrained rock data exist. These elements are calcium, potassium, CO2, and sulfur (Table 2). This is a somewhat random collection of elements and the comparisons thus derived may not be “typical” of all elemental fluxes, but they do give some first-order comparisons. The crustal fluxes of these elements add up to 3.6 £ 1014 g of seawater calcium, potassium, CO2, and sulfur per year, into the crust, while the sum of the flank and axis hydrothermal fluxes give a net flux out of the oceanic crust ranging from 1.8 £ 1014 g yr21 to 2.92 £ 1014 g yr21, depending on whether one uses the high or low estimates for axial fluxes. Overall, the rock record yields fluxes into the crust of the same order of magnitude but opposite sign to the fluid data, illustrating a rather fundamental disconnect between hydrothermal fluid data and ocean crust

alteration data. Ideally, fluxes derived from different and complementary data sets should be the same. What are the reasons for this discrepancy? Several contributing factors may be considered: (i) The two methods sample different types of processes. When Edmond et al. (1979) published the first “global” geochemical fluxes based on black-smoker data, Hart and Staudigel (1982) pointed out that there are major discrepancies with the fluxes of rubidium and caesium from oceancrust alteration data that appear to be largely controlled by low-temperature alteration. Fluxes based on hydrothermal fluids are biased towards high-temperature processes, which may underestimate the total flux. (ii) The two methods sample processes occurring at very different times in the hydrothermal history of the oceanic crust. The oldest hydrothermal vent samples analyzed to date come from Baby Bare seamount near the Juan de Fuca Ridge, on 3.5 Ma crust (Wheat and Mottl, 2000). At 3.5 Ma, the oceanic crust has lost only half of its total convective heat loss (C. A. Stein and S. Stein, 1994b). Hydrothermal fluxes from the rock record at sites 417 and 418, for example, offer a complete set of chemical changes occurring during the entire hydrothermal history of the oceanic crust. (iii) Oceanic crust has a substantial amount of pore space produced during its initial emplacement, and new pore space is generated as the crust cools. All these pores are filled with seawater and some additional water is taken up through hydration of ocean-crust materials. The chemical inventory of these pore waters are transferred into the crust, producing a one-way flux of water into the crust that is not accounted for through the measurement of hydrothermal fluids. This flux represents a baseline flux that should be subtracted from hydrothermal fluxes in order to calculate net fluxes out of the crust. It was estimated above that oceanic crust may take up to ,1.5 wt.% water. This translates into addition of ,2 £ 1012 g MgO yr 21, 0.5 £ 1012 g CaO yr 21, 1.5 £ 1013 g Na2O yr21, 1 £ 108 g Rb yr21, 0.3 £ 108 g CO2 yr21, and 1 £ 1013 g S yr21. Overall, these fluxes are relatively small when compared to the fluxes considered in this mass balance and are thus not likely to explain the discrepancy between hydrothermal- and ocean-crust-derived fluxes. (iv) Another issue in reconciling fluid versus crust fluxes relates to the nature of fluid circulation in the oceanic crust and the uncertainty in fluid pathways. With the exception of pore waters, fluids sampled from hydrothermal systems reflect relatively high water volumes, and high water– rock ratios. Most of the crust however, is altered at low water – rock ratios, in particular deep within the crust (.100m). Thus, it is possible that the alteration reflected in hydrothermal fluid data only

Discussion represents hydrothermal activity occurring in high water – rock ratio reaction zones and in high waterthroughput aquifers. By far the greatest volume of oceanic crust is altered at much lower water – rock ratios and by processes that may largely be diffusive. In addition, drilling recovery rates are particularly low in regions that are most porous, and thus high water –rock ratio regions, which are typically very porous, may be under-represented in crustal studies (v) The role of sediments in mass balances of seawater-ocean-crust chemical exchange remains largely unexplored. Pore-water studies in sediments on the ocean floor demonstrate that there is chemical exchange between sediments and the basaltic oceanic crust (Lawrence and Gieskes, 1981). The impact of these processes is not quantified but it is quite possible that this exchange has a significant affect on the bulk composition of the oceanic crust. This is particularly important for sedimented ridges and later stage alteration of the oceanic crust, when the crust is typically covered by sediments. Chemical exchange between sediment pore waters and the ocean crust may have a profound impact on the chemical fluxes into the oceanic crust, but it is not included in the hydrothermal spring data. At least four out of these five contributing factors may substantially contribute to the discrepancies between fluid- and crust-derived hydrothermal flux estimates.

3.15.5.1.3

Reconciling hydrothermal fluxes from fluid and rock data

The discrepancies discussed above suggest that both methods of flux determination may have intrinsic problems. Flux data from hydrothermal vents are compromised for determining global flux estimates because these fluids are derived from young crust, mostly from deeper crustal levels. They selectively sample high water – rock ratio alteration processes and they ignore lowtemperature fluid fluxes (,20 8C). Rock data are particularly unreliable or not available for the deeper crust and for high water – rock ratio reactions. Low recovery rates, and large primary variability in many important elements (magnesium, silicon) produce significant uncertainties in the flux estimates for elements whose fluxes are particularly well determined in hydrothermal fluids. The most important steps in reconciling these differences are to evaluate the weaknesses of both approaches and try to arrive at fluxes that use both methods in a complementary fashion. Furthermore, a critical error analysis will highlight the weaknesses of each of the methods, allowing the design of new experiments or approaches that resolve these issues by reducing

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a particular source of uncertainty, possibly separately for different element groups.

3.15.5.2 Impact of Ocean-crust Composition on Arc Processes and Mantle Heterogeneity Using average MORB or the range of compositions of oceanic basalts (e.g., Hofmann, 1988; Chapter 3.13 and http://petdb.ldeo.columbia.edu Lehnert et al., 2000), the fluxes derived here can be applied to determine the average compositions of oceanic crust that is subducted and recycled into the mantle. These compositions thus influence the composition of subduction zone magmas (see Chapter 3.18) and bear on the chemical mass balance of the mantle. Whereas average fluxes are useful for defining the global mass balance between mantle and the oceans, understanding the compositional diversity of subducted crust is important in constraining its dehydration or partial melting processes during recycling (see Chapter 3.17). Such diversity is reflected in the more extensively altered compositional domains in the oceanic crust such as volcaniclastics, and the moderately altered flow composites of sites 417A, 417D, and 418B (Staudigel et al., 1995, 1996). Other extreme compositional domains include umbers, ophicalcites (calcite-basalt breccia), massive sulfides, epidosites, talcum deposits, or serpentinites, but these are probably best studied in ophiolites. Particular mineralogical assemblages may show distinct phase relationships during prograde metamorphism. For example, ophicalcite or talcum has been shown to display distinct phase relations that is likely to control its dehydration or decarbonation behavior during subdution (Wyllie, 1978; Kerrick and Connelly, 2001). Volatile-rich or alkali-rich compositional domains may contribute preferentially to fluids that are extracted during prograde metamorphism in subduction zones, or they may melt prior to the surrounding less alkalic, and volatile-poor rock. The latter scenario has been considered in mantle melting models as a process that explains the isotopic variation in mantle-derived melts (Phipps-Morgan, 2001). MORB are generated by massive melting events that mix melts derived from depleted mantle and relatively enriched “plums” of recycled materials, while the smaller degrees of melting of ocean island basalts are capable of extracting more of these heterogeneities, displaying the fuller extent of mantle heterogeneity in bulk rock analyses. The geochemical characteristics of extremely altered compositional domains may help decipher the origin of volatiles or melts that may be related to the subduction of oceanic crust. Recycling of oceanic crust into the Earth’s mantle may profoundly influence the uranium

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budget, and the evolution of U/Th/Pb isotope systematics in the mantle (Hart and Staudigel, 1989; Elliott et al., 1999). Uranium is readily taken up by oceanic crust during hydrothermal alteration and recycled into the mantle while thorium concentrations remain relatively unchanged during ocean-crust alteration. The uptake of uranium in the oceanic crust is restricted to the upper alteration zones, having relatively high water –rock ratios. There, oxidizing seawater enters the crust and loses its dissolved oxygen from reactions with the highly reducing oceanic crust. Most of the uranium dissolved in seawater is 6þ in the oxidized form of UO2(CO3)22 2 (i.e., U ), 4þ and is reduced to U , which is as insoluble in hydrous solutions as thorium (Langmuir, 1978). This fixation of uranium in reducing hydrothermal environments is critically dependent on its mobilization by oxidation under the present-day atmospheric conditions. However, the earth’s atmosphere has only been oxidizing for the last 2.2 Gyr (Holland, 1984, 1994), and for this reason, uranium recycling is unlikely to have occurred prior to 2.2 Ga. This process may explain the “kappa-conundrum,” whereby MORB and the upper mantle appears to have much lower 232Th/238U ratios ( ¼ “kappa”) than required by modeling of lead isotope ratios (Elliott et al., 1999 and references therein). These present-day low kappas may be caused by the recycling of uranium relatively recently in Earth’s history. One of the major fluxes associated with the recycling of oceanic crust involves water and CO2. Most of the volatile inventory of altered oceanic crust is located in the uppermost 600 m, which is also the section first exposed to the top-down heating of the slab during subduction. For this reason, extraction of volatiles during subduction is particularly efficient in the upper part of the slab and much of this inventory is likely to be extracted (e.g., Kerrick and Connolly, 2001). However, there are several mechanisms that could allow these elements to survive passage through the “subduction zone filter (see also Chapter 3.17).” (i) In particular old and dense crust may be subducted relatively rapidly, which greatly reduces the geotherms in the subducting slab, increasing volatile subduction (Staudigel and King, 1992). (ii) Uneven topography on the surface of subducting upper oceanic crust (due to horst-graben structures and collapsed seamounts) can produce vertical throw of highly altered materials to levels substantially below the average top of the remaining slab, thus isolating these materials from top-down heating. (iii) Some volatiles from dehydration may be retained as fluid inclusions in newly formed minerals that remain stable past the subduction zone, and may make up several percent of a rock (Touret and Olsen, 1985).

3.15.6 CONCLUSIONS Seafloor hydrothermal alteration processes are important for the global geochemical cycles of many elements, and the record of these processes in the oceanic crust reveals much information about these cycles. Rather robust flux information can be obtained from a variety of elements that have rather low initial abundances in basalt (H2O, CO2, K2O, rubidium, caesium, uranium) or that are rather sensitive to alteration (87Sr/86Sr) and d18O. Fluxes of many other elements are rather poorly constrained because of substantial primary magmatic variation. This rock record yields results that are inconsistent with fluxes inferred from fluid data from seafloor hydrothermal springs. An in-depth analysis of data methods and uncertainties in fluxes based on rock data and fluid data suggest that these discrepancies are due to profound problems with both types of flux determination. Fluxes from fluid data are fundamentally limited by their near-exclusive focus on reactions involved in large fluid fluxes, hightemperature reactions and in particular in young crust near the ridge axis. A few studies have recently begun exploring low-temperature vents in crust up to 3.5 Myr old. Hydrothermal fluxes up to this age involves only about half of the total convective heat flow lost in the oceans, and by far the largest volume of fluids passes through oceanic crust older than 3.5 Ma. Furthermore, much, if not most of the oceanic crust is altered at low fluid/rock ratios, and thus, most of the alteration of the oceanic crust is not accounted for by in these fluxes. Data from the rock record are limited to the rather small number of drill sites that are in crust old enough and have sufficiently high recovery rates to allow reliable estimates to be made. None of the drill holes available so far reaches into the reaction zone of a black smoker, which is where much of the hydrothermal flux data derive from. Most importantly, the uncertainties of some element fluxes (i.e., silicon, aluminum, magnesium, etc.) derived from oceanic crust studies are substantially larger than the fluxes expected from balancing other flux data. This is due to the rather high concentrations of these elements in the oceanic crust, their large magmatic variation, and their complex behavior during alteration. It is thus unlikely that the current approach will yield fluxes for these elements that are sufficiently constrained to be meaningful in the context of global fluxes in the hydrosphere. There is no simple solution to “fix” the intrinsic problems of fluid or rock-based estimates of hydrothermal fluid fluxes. For this reason, both methods have to be used in a complementary fashion. The first goal in such a complementary

References analysis is to develop a reference model for oceancrust hydrothermal alteration, with a clear definition of reservoirs, reaction zones, types of alteration, etc. The goal of such a reference model is to define the portions of oceanic crust that can be constrained by various methods. Fluxes derived from studies of ophiolites need to be used to constrain those of the deeper oceanic crust, until all major fractions of the oceanic crust can be reliably recovered by drilling. Above all, studies should evaluate uncertainties and place all flux determinations in this context.

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Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 511–535

3.16 Oceanic Plateaus A. C. Kerr Cardiff University, Wales, UK 3.16.1 INTRODUCTION 3.16.2 FORMATION OF OCEANIC PLATEAUS 3.16.3 PRESERVATION OF OCEANIC PLATEAUS 3.16.4 GEOCHEMISTRY OF CRETACEOUS OCEANIC PLATEAUS 3.16.4.1 General Chemical Characteristics 3.16.4.2 Mantle Plume Source Regions of Oceanic Plateaus 3.16.4.3 Caribbean–Colombian Oceanic Plateau ( ,90 Ma) 3.16.4.4 Ontong Java Plateau ( ,122 and ,90 Ma) 3.16.5 THE INFLUENCE OF CONTINENTAL CRUST ON OCEANIC PLATEAUS 3.16.5.1 The North Atlantic Igneous Province ( ,60 Ma to Present Day) 3.16.5.2 The Kerguelen Igneous Province ( ,133 Ma to Present Day) 3.16.6 IDENTIFICATION OF OCEANIC PLATEAUS IN THE GEOLOGICAL RECORD 3.16.6.1 Diagnostic Features of Oceanic Plateaus 3.16.6.2 Mafic Triassic Accreted Terranes in the North American Cordillera 3.16.6.3 Carboniferous to Cretaceous Accreted Oceanic Plateaus in Japan 3.16.7 PRECAMBRIAN OCEANIC PLATEAUS 3.16.8 ENVIRONMENTAL IMPACT OF OCEANIC PLATEAU FORMATION 3.16.8.1 Cenomanian–Turonian Boundary (CTB) Extinction Event 3.16.8.2 Links between CTB Oceanic Plateau Volcanism and Environmental Perturbation 3.16.9 CONCLUDING STATEMENTS REFERENCES

537 539 540 540 540 541 544 548 549 549 550 551 552 553 554 556 557 558 558 560 561

knowledge of the ocean basins has improved over the last 25 years, many more oceanic plateaus have been identified (Figure 1). Coffin and Eldholm (1992) introduced the term “large igneous provinces” (LIPs) as a generic term encompassing oceanic plateaus, continental flood basalt provinces, and those provinces which form at the continent – ocean boundary (volcanic rifted margins). LIPs are generally believed to be formed by decompression melting of upwelling hotter mantle, known as mantle plumes. Although ideas about hotpots and mantle plumes have been around for almost 40 years (Wilson, 1963), it is only in the past 15 years that LIPs have become the focus of major research. One of the main reasons for the increased research activity into LIPs is the realization that significant proportions of these LIPs erupted over a relatively short time, often less than 2– 3 Myr (see review in Coffin, 1994).

3.16.1 INTRODUCTION Although the existence of large continental flood basalt provinces has been known for some considerable time, e.g., Holmes (1918), the recognition that similar flood basalt provinces also exist below the oceans is relatively recent. In the early 1970s increasing amounts of evidence from seismic reflection and refraction studies revealed that the crust in several large portions of the ocean floor is significantly thicker than “normal” oceanic crust, which is 6 –7 km thick. One of the first areas of such over-thickened crust to be identified was the Caribbean plate (Edgar et al., 1971) which Donnelly (1973) proposed to be an “oceanic flood basalt province”. The term oceanic plateau was coined by Kroenke (1974), and was prompted by the discovery of a large area of thickened crust (.30 km) in the western Pacific known as the Ontong Java plateau (OJP). As our 537

538

Oceanic Plateaus

This has important implications for mantle processes and source regions (Hart et al., 1992; Stein and Hofmann, 1994), as well as environmental effects on the global biosphere (e.g., Caldeira and Rampino, 1990; Courtillot et al., 1996; Kerr, 1998). Oceanic plateaus can also become accreted to continental margins, and it has been proposed that these plateaus have been significant contributors to the growth of continental crust (e.g., Abbott, 1996; Albarede, 1998). The most recent major phase of oceanic plateau formation was in the Cretaceous when the Ontong Java, Manihiki, Hess Rise, and the Caribbean –Colombian plateaus formed in the

Pacific, while in the Indian Ocean the Kerguelen plateau was developing. The areas, volume maximum thicknesses and ages of the larger of these plateaus are given in Table 1. The Ontong Java is the largest of the Cretaceous plateaus. It covers an area of 1.9 £ 106 km2, and has an estimated total volume of 4.4 £ 107 km3 (Eldholm and Coffin, 2000). Although early seismic refraction data suggested that the OJP was as thick as 43 km (Furomoto et al., 1976), a more recent synthesis based on existing seismic and new gravity data (Gladczenko et al., 1997) has indicated the average thickness to be ,32 km.

Figure 1 Map showing all major oceanic plateaus, and other large igneous provinces discussed in the text (after Saunders et al., 1992). Table 1 Ages and dimensions of Jurassic– Cretaceous oceanic plateaus. Oceanic plateau Hikurangi Shatsky Rise Magellan Rise Manihiki Ontong Java Hess Rise Caribbean South Kerguelen Central Kerguelen/Broken Ridge Sierra Leone Rise Maud Rise After Eldholm and Coffin (2000).

Mean age (Ma)

Area (106 km2)

Thickness range (km)

Volume (106 km3)

early-mid Cretaceous 147 145 123 121(90) 99 88 110 86 ,73 ,. 73

0.7 0.2 0.5 0.8 1.9 0.8 1.1 1.0 1.0 0.9 0.2

10– 15 10– 28 10 .20 15– 32 .15 8 – 20 ,22 19– 21 .10 .10

2.7 2.5 1.8 8.8 44.4 9.1 4.4 6.0 9.1 2.5 1.2

Formation of Oceanic Plateaus

539

3.16.2 FORMATION OF OCEANIC PLATEAUS The production of large volumes (.106 km3) of melt in a period as short as 2 – 3 Myr implies magma production rates up to 25% higher than those observed at present-day midocean ridges (Eldholm and Coffin, 2000), and is generally believed to necessitate a high flux of hotter-thanambient asthenospheric mantle below these provinces (e.g., McKenzie and Bickle, 1988). Numerical and physical models show that this hotter mantle commonly takes the form of a mantle plume which ascends by thermal buoyancy through the overlying mantle (Loper, 1983; McKenzie and Bickle, 1988; Campbell et al., 1989; Farnetani and Richards, 1995). Physical constraints demand that mantle plumes must ascend from a boundary layer within the Earth, either the core –mantle boundary (D00 ) or the 670 km discontinuity. Large ascending mantle plumes are, on average, 200 8C hotter than the ambient upper mantle (McKenzie and Bickle, 1988) and undergo decompression melting as they approach the base of the lithosphere. Physical modeling experiments by Griffiths and Campbell (1990) have shown that mantle plumes are likely to ascend through the mantle from their source boundary layer in the form of a large semi-spherical “head” fed from the source region by a narrower plume tail (Figure 2). Alternatively, numerical modeling by Farnetani and Richards (1995) suggested that plume heads starting in the mantle only rise about three plume head diameters before spreading out. In either case, as the plume approaches the base of the lithosphere, it spreads out over a broadly circular area (which can be as much as 1000 km in diameter) and undergoes adiabatic decompression, producing melt over most of the area covered by the flattened-out plume head (Campbell and Griffiths, 1990). The amount of melt produced is critically dependent on the thickness of the preexisting lithosphere, since the base of the rigid, nonconvecting lithosphere will act as a “lid” on the upwelling plume mantle and on the extent of decompression melting. Thus, a mantle plume ascending below thick continental lithosphere (.50 km) will produce a smaller thickness of melt than a plume which ascends beneath oceanic lithosphere (#7 km) (Figure 3). Another significant factor in determining the amount of melt generated by a mantle plume is the temperature of the plume: generally the higher the temperature, the more melt will be produced (Figure 3). The initial 40Ar/39Ar step-heating ages for LIPs support models of rapid formation and eruption, often in less than 2– 3 Myr (Richards et al., 1989). As more age data have become available, a wider age-range has emerged for some LIPs (e.g., the

Figure 2 Cartoon to illustrate how mantle plumes are believed to (a) rise through the asthenosphere from either the 670 km discontinuity or D00 and (b) flatten along the base of the lithosphere and undergo decompression melting to produce a LIP (after Saunders et al., 1992).

Figure 3 Schematic diagram showing how original lithospheric thickness and mantle potential temperature affect the amount of melt produced (melt thickness) and how these factors relate to continental flood basalts (CFB), volcanic rifted margins (VRM), off – ridge and ridge – centered oceanic plateaus (OP), and midocean ridges (MOR).

Caribbean plateau, 95 – 86 Ma: Kerr et al., 1997a; Sinton et al., 1998; Hauff et al., 2000b). Nevertheless, it still appears that substantial proportions of these provinces were formed over geologically

540

Oceanic Plateaus

short time periods (e.g., the Ontong Java plateau formed on two occasions: 122^3 and 90^4 Ma; Neal et al., 1997). Theoretically, oceanic plateaus can form anywhere in the ocean basins; however, most oceanic plateaus appear to have formed at or near midocean ridges (e.g., Kerguelen, Manihiki and Ontong Java), i.e., regions that are conducive to voluminous decompression melting (Eldholm and Coffin, 2000). At first glance, it appears somewhat coincidental that plumes of deep mantle origin reach the base of the lithosphere at a midocean ridge. However, as pointed out by Saunders et al. (1996), a likely explanation for this observation is that mantle plumes can “capture” oceanic spreading centers (cf. present-day Iceland). 3.16.3 PRESERVATION OF OCEANIC PLATEAUS The oldest in situ oceanic crust is Jurassic in age (Pringle, 1992), because within 200 Myr of its formation at a midocean ridge, the oceanic crust was recycled back into the asthenosphere. Although many of the Cretaceous and late Jurassic oceanic plateaus still form part of the ocean basins, the preservation potential of oceanic plateaus older than Jurassic is low. Fortunately, however, oceanic plateaus are much more buoyant than oceanic crust of “normal” thickness formed at a midocean ridge (e.g., Cloos, 1993). This excess buoyancy is primarily due to the greater crustal thickness of oceanic plateaus if the plateaus are relatively young; residual heat from their formation can also contribute to their buoyancy. Recent measurements of upper mantle shear wave splitting and shear wave velocity structure (Klosko et al., 2001) reveal that the Ontong Java plateau is underlain by a 300 km thick, long-lived, rheologically strong and chemically depleted root. Klosko et al. (2001) propose that this root represents the residue from mantle melting and that its consequent lower density contributes significantly to the buoyancy of both the Ontong Java and other oceanic plateaus. The net result of this excess buoyancy is that oceanic plateaus, in contrast to normal oceanic crust, are much less easily subducted (Ben-Avraham et al., 1981; Cloos, 1993; Kimura and Ludden, 1995). Instead of being completely recycled back into the mantle, their upper layers can be “peeled off” (Kimura and Ludden, 1995) and accreted on to the margin of the subduction zone. This plateau accretion can occur either at an Andean-type continental margin, or an island arc. Although many of the in situ Cretaceous oceanic plateaus have been drilled by the Deep Sea Drilling (DSDP) and Ocean Drilling Programs (ODP), the insight that these drill holes can provide is relatively limited compared to the

accreted oceanic plateau sections. For example, the Ontong Java plateau collided with the westward-dipping Solomon Islands subduction zone at 10– 20 Ma, resulting in a reversal in the polarity of subduction from west to east, and the uplift and exposure of the deeper sections of the plateau on the Solomon Islands (Neal et al., 1997; Petterson et al., 1999). A second example is the Caribbean – Colombian oceanic plateau, which formed in the Pacific at ,90 Ma (Sinton et al., 1998; Hauff et al., 2000a). Within 10 Myr the eastward-moving Farallon plate had brought the southern portion of the plateau into collision with continental northwestern South America, resulting in the accretion of slices of the plateau on to the continental margin (Kerr et al., 1997b). This accretion was accompanied by back-stepping of the subduction zone west of the accreted plateau slices. Shortly after its formation the northern portion of the plateau began to move into the proto-Caribbean seaway between North and South America (Burke, 1988; Kerr et al., 1999; White et al., 1999). In doing so the plateau encountered the eastward-dipping “Great Arc of the Caribbean”. Unable to subduct, the thick plateau clogged the subduction zone, resulting in a reversal in the polarity of subduction from east to west. This reversal in subduction polarity in conjunction with a back-stepping of subduction behind the advancing plateau (Burke, 1988), effectively isolated the Caribbean as a separate plate. Over the past ,80 Ma the northern portion of the Caribbean plateau has moved into the gap between North and South America and has been uplifted and subaerially exposed around its margins (Figure 4), thus making it available for detailed study. It is in ways such as these that remnants of these thick, buoyant oceanic plateaus can be preserved and incorporated into the continental crust. The identification of these older preserved plateaus within the geological record will be explored in a later section.

3.16.4 GEOCHEMISTRY OF CRETACEOUS OCEANIC PLATEAUS 3.16.4.1 General Chemical Characteristics Table 2 shows representative analyses of in situ and accreted portions of Cretaceous oceanic plateaus. Cretaceous oceanic plateaus are predominantly basaltic (,10 wt.% MgO) in composition (Figure 5), although more-MgO-rich lava flows are found in the Caribbean – Colombian oceanic plateau (CCOP). Typically, oceanic plateaus possess generally low levels of compatible elements (i.e., Ni , 300 ppm and Cr , 1,000 ppm; see Figure 6). In terms of incompatible trace elements the majority of oceanic plateau lavas and intrusive sheets possess relatively flat rare earth

Geochemistry of Cretaceous Oceanic Plateaus

541

Figure 4 Map to show the main accreted outcrops of the Caribbean – Colombian oceanic plateau along with the locations of DSDP/ODP drill holes which penetrated the thickened crust of the Caribbean plate.

primitive mantle-normalized patterns with abundances varying between 5 and 10 times primitive mantle values (Figure 7). The radiogenic isotope compositions of oceanic plateaus have been well characterized: initial 1Nd values for oceanic plateaus generally range from þ6.0 to þ9.0, whereas initial 87Sr/86Sr ratios fall mostly between 0.703 and 0.704 (Figure 8), i.e., the 87Sr/86Sr and 1 . Nd values are generally less depleted in terms of their radiogenic isotopes than “normal” (N)-MORB as typified by MORB from the East Pacific Rise (Figure 8(a)). Elevated initial 87 Sr/86Sr ratios (.0.7040) are most likely due to secondary alteration by hydrothermal fluids. It is noteworthy that the high-MgO rocks generally possess more extreme (1Nd)i values than the basalts (Figure 8(a)). Figure 8(b), a plot of 207 Pb/204Pb against 206Pb/204Pb, reveals that most oceanic plateau rocks range between 18.5 and 19.5 206Pb/204Pb and between 15.525 and 15.625 207Pb/204Pb. An interesting feature of Figure 8(b) is that many of the basalts lie on a trend between the proposed mantle components of HIMU and DMM (Zindler and Hart, 1986) while the high-MgO lavas form a trend between DMM and the enriched mantle component EM2. Zindler and Hart (1986) proposed four mantle components: depleted MORB mantle (DMM), two

types of enriched mantle EM1 and EM2, and HIMU, so-called because it has a high 238U/204Pb ratio, or m value. In recent years our knowledge of the radiogenic isotope systematics of oceanic plateaus has been augmented by the analysis of Hf and Os isotopes. Although data for the Kerguelen and Ontong Java plateaus are still relatively sparse, more data exist for the CCOP. Initial 1Hf values for the CCOP and OJP range from 10 to 18 (Figure 9(a)). The CCOP samples form a trend between MORB source mantle and the HIMU component, while the OJP samples possess lower initial 1Nd values at equivalent 1Hf than the CCOP and appear to form a trend towards EM2. The much lower 1Hf and 1Nd for the Kerguelen plateau (Figure 9(a)) will be discussed in Section 3.16.5.2. Initial g Os for high-MgO rocks from the CCOP range from 0 to þ18, whereas the basalts range from 27 to þ10 (Figure 9(b)). How representative these ranges are for g Os, requires the acquisition of more data from other oceanic plateaus. 3.16.4.2 Mantle Plume Source Regions of Oceanic Plateaus Since the pioneering study of Hoffman and White, it has become widely accepted that one of

Table 2 Representative analyses of Cretaceous oceanic plateau lavas. Plateau

Location

Sample

Data sources

Zr

Kerguelen Kerguelen Kerguelen Kerguelen Kerguelen Ontong Java Ontong Java Ontong Java Ontong Java Ontong Java CCOP CCOP CCOP CCOP CCOP CCOP CCOP CCOP CCOP CCOP

ODP site 747 ODP site 748 ODP site 749 ODP site 750 ODP site 738 ODP site 807 ODP site 807 Santa Isabel Maliata Maliata Gorgona Gorgona Gorgona Colombia Colombia Colombia Curac¸ao Curac¸ao DSDP site 150 Ecuador

16-5,103-6 79-6,90-4 15-5,125-7 17-3,23-26 34-1,88-92 75-4,46-48 88-3,76-79 I96 SG1 ML407 GOR160 GOR117 GOR94-35 SDB18 VIJ1 COL472 CUR14 CUR20 11-2, 63-67 EQ1

1 1 1 1 2 3 3 4 5 5 6 6 7 8 8 9 10 10 11 12

50.90 49.24 52.73 49.21 52.17 48.75 49.74 49.28 50.38 49.67 45.13 50.40 46.58 50.63 51.45 48.81 46.42 52.13 49.46 48.88

Kerguelen Kerguelen Kerguelen Kerguelen Kerguelen Ontong Java Ontong Java Ontong Java Ontong Java Ontong Java CCOP CCOP CCOP CCOP CCOP CCOP

ODP site 747 ODP site 748 ODP site 749 ODP site 750 ODP site 738 ODP site 807 ODP site 807 Santa Isabel Maliata Maliata Gorgona Gorgona Gorgona Colombia Colombia Colombia

16-5,103-6 79-6,90-4 15-5,125-7 17-3,23-26 34-1,88-92 75-4,46-48 88-3,76-79 I96 SG1 ML407 GOR160 GOR117 GOR94-35 SDB18 VIJ1 COL472

1 1 1 1 2 3 3 4 5 5 6 6 7 8 8 9

Rb 18.5 7.8 12.3 9.0 37.4 1.0 10.0 2.0 1.3 1.8 1.0 0.3 4.2 2.5 0.4 1.9

Nb

Mo

Hf

Pb

Ta

1.21 2.74 1.52 1.17 1.73 1.62 1.13 1.31 1.65 1.35 0.64 0.77 0.37 1.47 2.02 0.89 0.57 0.78 1.27 1.23

17.23 17.50 15.18 15.77 15.44 14.16 14.27 14.20 13.54 13.64 11.78 13.81 10.57 14.19 12.85 12.86 9.53 12.85 16.60 14.39

10.62 8.77 11.95 14.62 10.82 13.43 13.44 12.44 14.01 13.47 12.52 12.62 11.41 13.34 15.20 10.66 11.06 9.80 10.30 12.75

0.16 0.09 0.17 0.18 0.18 0.21 0.22 0.16 0.19 0.19 0.18 0.23 0.17 0.23 0.19 0.17 0.17 0.20 0.10 0.20

7.47 7.54 7.22 8.36 5.51 6.74 7.43 8.03 7.30 7.45 18.25 8.65 21.96 6.77 5.78 11.55 22.86 8.60 8.58 8.96

Cs ,0.8 2.81 ,0.5 ,0.9

0.01 0.01

Sr 234 1131 214 193 273 174 107 115 108 100 64 107 34 156 89 398

Ba 229 1661 114 30 336

37 28 23 8 21 20 37 27 85

Sc 33.8 18.7 34.5 38.9 36.9 45.9 52.3 50.0 41.0 40.0 28.3 41.1 27.3 42.2 44.4 30.8

V 201 170 271 269 267 313 349 341 392 295 227 371 166 331 531 334

Th

U

9.59 6.71 8.31 9.01 9.17 12.13 12.00 12.50 11.64 11.38 10.10 12.39 8.42 10.94 9.57 11.61 8.20 14.30 9.87 10.81

1.88 4.72 3.42 1.99 2.98 2.38 2.06 1.89 2.32 3.08 1.30 2.43 0.55 2.89 2.29 2.91 0.90 1.43 2.60 1.89

Cr 401 166 260 193 95 162 163 238 56 61 1373 194 82 208 63 1393

Ni 72 182 30 120 28 99 87 122 61 62 723 112 968 97 58 264

87

Sr/ 86Sri

2.00 0.56 0.54 0.19 1.95 0.13 0.28 0.15 0.15 0.26 0.06 0.02 0.19 0.15 0.07 0.08 0.02 0.03 0.13 0.72

1Ndi

Pb/ 204Pb

207

Pb/ 204Pb

101.22 98.99 101.22 100.61 100.16 99.69 100.66 100.06 101.32 100.62 100.01 101.38 100.25 100.72 99.59 99.64 99.79 100.18 99.02 100.01

3.63 7.14 1.35 4.31 0.58 0.45 20.24 1.61 0.30 2.14

Zn 66 79 115 93 102

Ga 22.0 18.7 22.7 19.8 20.7

124 229

93 57

2030

56

133 125

124 70

17.0 17.0 13.2 15.7 10.8 19.1 20.2 14.3

La 12.30 105.00 6.80 4.00 17.10 6.15 2.96 3.40 4.75 3.86 0.65 1.02 0.22 3.80 6.05 2.42

Cu

0.16 1.12 0.18 0.11 0.21 0.14 0.09 0.10 0.14 0.13 0.05 0.06 0.03 0.12 0.17 0.09 0.05 0.06 0.10 0.18

206

208

Pb/ 204Pb

6.0 5.0

3.26 1.57 0.90 3.03 2.23 1.41 4.45 Ce 25.50 224.00 16.10 8.90 39.30 14.59 8.07 9.90 12.60 10.60 2.17 2.99 0.57 9.80 16.55 6.33

CCOP CCOP CCOP CCOP

Curac¸ao Curac¸ao DSDP site 150 Ecuador

CUR14 CUR20 11-2, 63-67 EQ1

10 10 11 12

Kerguelen Kerguelen Kerguelen Kerguelen Kerguelen Ontong Java Ontong Java Ontong Java Ontong Java Ontong Java CCOP CCOP CCOP CCOP CCOP CCOP CCOP CCOP CCOP CCOP

ODP site 747 ODP site 748 ODP site 749 ODP site 750 ODP site 738 ODP site 807 ODP site 807 Santa Isabel Maliata Maliata Gorgona Gorgona Gorgona Colombia Colombia Colombia Curac¸ao Curac¸ao DSDP site 150 Ecuador

16-5,103-6 79-6,90-4 15-5,125-7 17-3,23-26 34-1,88-92 75-4,46-48 88-3,76-79 I96 SG1 ML407 GOR160 GOR117 GOR94-35 SDB18 VIJ1 COL472 CUR14 CUR20 11-2, 63-67 EQ1

1 1 1 1 2 3 3 4 5 5 6 6 7 8 8 9 10 10 11 12

Kerguelen Kerguelen Kerguelen Kerguelen Kerguelen Ontong Java Ontong Java Ontong Java Ontong Java Ontong Java CCOP

ODP site 747 ODP site 748 ODP site 749 ODP site 750 ODP site 738 ODP site 807 ODP site 807 Santa Isabel Maliata Maliata Gorgona

16-5,103-6 79-6,90-4 15-5,125-7 17-3,23-26 34-1,88-92 75-4,46-48 88-3,76-79 I96 SG1 ML407 GOR160

1 1 1 1 2 3 3 4 5 5 6

1.7 0.4 3.2 5.2 Pr

1.56 1.98 1.70 0.46 0.60 0.13

0.92 0.79 1.36 1.56 Zr 97 599 91 47 166 98 64 65 81 83 29

46 69 117 111 Nd 13.50 103.00 11.90 6.30 21.20 11.65 6.50 7.84 4.61 5.67 2.86 3.53 0.90 9.30 12.0 5.58 3.16 4.80 6.94 8.11

Sm 3.50 14.20 3.64 2.32 5.28 3.39 2.30 2.77 1.69 2.11 1.25 1.74 0.65 2.85 3.97 1.76

Nb 7.00 121.90 5.80 3.33 8.90 5.70 3.32 4.30 4.90 5.00 0.48

Mo

1.39 2.32 2.57

0.59 0.36

6 11 16 15 Eu 1.03 3.31 1.24 0.84 1.73 1.31 0.93 1.02 1.13 0.94 0.48 0.63 0.34 1.11 1.47 0.74 0.38 0.55 0.89 0.96 Hf 2.34 12.40 2.16 1.16 3.89 2.64 1.74 2.50 2.63 1.00

29.7 43.3 61.5 Gd 0.00 8.90 3.00 4.51 3.39 3.50 4.14 3.54 1.58 2.07 1.50 2.94 4.76 2.67 1.87 2.87 3.46 Pb 8.00 1.63 0.82

0.23 0.17 0.15 1.30

186 258 335 353 Tb 0.65 1.25 0.69 0.58 0.94 0.81 0.70 0.63 0.73 0.68 0.29 0.41 0.30 0.80 0.93

0.52 0.66 Ta 0.46 9.08 0.40 0.19 0.54 0.34 0.20 0.35 0.25

2017 552 373 285 Dy

1032 178 127 101

75 102 150

Ho

Er

4.20 4.56 4.22 2.10 3.13 2.18

0.91 1.00 0.92

0.48

2.64 2.90 2.57 1.43 2.18 1.44

2.92

0.58

1.44

2.11 3.36 4.40

0.45 0.69 0.90

1.38 1.89 2.60

Th 0.20 13.20 1.80 1.40 2.27 0.56 0.34 0.27 0.28 0.16 0.02

U 0.000 2.200 0.148 0.048

0.120 0.097 0.047 0.011

87

Sr/86Sri 0.705783 0.705319 0.704260 0.706165 0.709730 0.704330 0.703560 0.703690 0.704040 0.704130 0.703041

70 66 85

9.1 12.2 17.0

1.20 2.84 3.16 3.50

Tm

Yb 1.89 1.80 2.71 2.57 2.87 2.71 2.40 2.30 2.83 2.47 1.30 2.20 1.42 2.75 4.25 1.36 0.50 1.36 1.82 2.53

Lu 0.28 0.26 0.39 0.34 0.44 0.42 0.38 0.35 0.43 0.37 0.21 0.32 0.21 0.43 0.63 0.21

0.37 0.41 0.36 0.22 0.37 0.23

0.28 0.39 1Ndi 24.0 23.3 1.8 1.4 28.2 5.0 5.9 5.4 3.9 5.7 9.5

206

Pb/204Pb 17.65 18.19 18.03 17.53 17.82 18.40 18.67 18.64 17.85 18.32 18.32

4.20 7.59 8.55 9.63 Y 23.2 27.7 29.3 24.6 28.8 30.0 24.0 22.0 25.0 23.0 13.8 22.8 15.3 29.4 41.7 19.0 11.5 15.8 21.0 22.3

0.21 0.26 0.38 207

Pb/204Pb 15.51 15.64 15.55 15.49 15.75 15.53 15.55 15.54 15.47 15.51 15.50

208

Pb/204Pb 38.16 38.38 38.16 38.01 39.01 38.38 38.54 38.60 37.91 38.23 37.82

(continued)

0.094 0.080 0.23 0.58

0.10 0.90 0.26 0.25

0.210

1.92 1.89

2.37 1.20 1.00 0.23 0.14

0.30 0.48 2.05 3.14

Sources: 1 Salters et al. (1992), 2 Mahoney et al. (1995), 3 Mahoney et al. (1993a), 4 Tejada et al. (1996), 5 Tejada et al. (2002), 6 Aitken and Echeverrı´a (1984), Dupre´ and Echeverrı´a (1984), Jochum et al. (1991), 7 AC Kerr unpublished data, 8 Kerr et al. (1997), Hauff et al. (2000b), 9 Kerr et al. (2002), 10 Kerr et al. (1996b), 11 Hauff et al. (2000b), 12 Reynaud et al. (1999).

15.58 15.54 15.54 15.58 15.58 15.59 15.55 15.60 15.53 18.86 18.68 18.87 19.22 19.33 19.31 19.08 19.07 18.16 8.6 9.4 8.5 8.1 11.0 6.6 7.0 7.2 10.1 0.703283 0.704767 0.703380 0.703207 0.513197 0.702961 0.703215 0.703546 0.703200 0.034

0.08 1.20 0.30 0.62 12.10 1.05

0.87 0.80 4.29 8.00 2.16 2.50 3.40 11.00 4.16 39 13 70 119 40 30 40 68 60 6 7 8 8 9 10 10 11 12 Gorgona Gorgona Colombia Colombia Colombia Curac¸ao Curac¸ao DSDP site 150 Ecuador CCOP CCOP CCOP CCOP CCOP CCOP CCOP CCOP CCOP

GOR117 GOR94-35 SDB18 VIJ1 COL472 CUR14 CUR20 11-2, 63-67 EQ1

Location Plateau

Table 2

(continued).

Sample

Data sources

Zr

Nb

Mo

Hf

Pb

Ta

Th

U

87

Sr/ 86Sri

1Ndi

206

Pb/ 204Pb

207

Pb/ 204Pb

208

38.56 38.19 38.46 38.91 38.86 38.90 38.82 38.70 37.84

Oceanic Plateaus Pb/ 204Pb

544

the principal contributors to the source regions of deep mantle plumes are subducted oceanic slabs and their sediments which descended through the asthenospheric mantle and ponded at either the 670 km discontinuity or the core –mantle boundary (D00 ). Zindler and Hart (1986) identified three main mantle plume components: HIMU (proposed as being derived from subducted oceanic crust); and two enriched components, EM1 and EM2. In addition, the upper asthenosphere was proposed as consisting of depleted MORB mantle (DMM). More recently, it has been shown that depleted signatures found in some LIPs represent a component which, rather than being due to entrainment of depleted upper mantle material (DMM), is derived from depth, and so is an integral part the plume itself (Kerr et al., 1995a,b; Fitton et al., 1997; Kempton et al., 2000). Kerr et al. (1995a,b) and Walker et al. (1999) proposed that this depleted plume component was ultimately derived from subducted oceanic lithosphere, unlike HIMU, which has its source in more-enriched upper oceanic crust. Evidence for recycled oceanic crust and lithosphere in the mantle plume source regions of oceanic plateaus has been presented by several authors (e.g., Walker et al., 1999; Hauff et al., 2000a), and mixing trends between a depleted component and HIMU are clearly seen on most of the radiogenic isotope plots (Figures 8 and 9).

3.16.4.3 Caribbean–Colombian Oceanic Plateau (,90 Ma) The Caribbean – Colombian Oceanic Plateau (CCOP) is exposed around the margins of the Caribbean and along the northwestern continental margin of South America (Figure 4). The thickened nature of the bulk of the Caribbean plate (8 –20 km; Edgar et al., 1971; Mauffret and Leroy, 1997) testifies to its origin as an oceanic plateau. The plateau has been drilled by DSDP Leg 15 and ODP Leg 165 (Figure 4; Bence et al., 1975; Sinton et al., 2000). The accreted plateau material in Colombia, Ecuador, Costa Rica and Hispaniola consists of fault-bounded slices of basaltic, and occasionally picritic lavas and sills with relatively few intercalated sediments and ash layers (Kerr et al., 1997a). Although they preserve layered and isotropic gabbros and ultramafic rocks, unlike accreted ophiolites generated at spreading centers, these accreted sequences of oceanic plateau do not possess sheeted dyke complexes (Kerr et al., 1998). Several other exposures are worthy of special mention: firstly, the 5 km thick section on the island of Curac¸ao, 70 km north of the coast of Venezuela (Figure 4). The sequence consists of

Geochemistry of Cretaceous Oceanic Plateaus

545

Figure 5 Plots of major elements (wt.%) against wt.% MgO for lavas from throughout the CCOP. All lavas with .18 wt.% MgO contain accumulated olivine. Data sources: Colombia—Kerr et al. (1997a), Gorgona—Echeverrı´a (1980), Kerr et al. (1996a), Arndt et al. (1997), Costa Rica—Hauff et al. (2000b), Curac¸ao—Kerr et al. (1996b). Circled fields are from the Ontong Java plateau (Mahoney et al., 1993a,b). Where the A and CG types differ markedly in composition they are plotted as separate fields, otherwise the OJP is plotted as a single field. OJP Site 1185 unpublished data were kindly provided by Godfrey Fitton.

pillowed picrites (MgO .12 wt.%) low in the succession that gradually give way to more basaltic pillow lavas nearer the top. These pillow lavas are intercalated with hyaloclastite horizons and intrusive sheets (Klaver, 1987; Kerr et al., 1996b). The second noteworthy locality is the island of Gorgona, 50 km off the western coast of Colombia (Figure 4). This small island (2.5 £ 8 km) is the site of the youngest komatiites (MgO-rich lava flows: .15 wt.%), which possess platy and blade-shaped olivines, giving the rocks a characteristic “spinifex” texture (Echeverrı´a, 1980; Kerr et al., 1996a). Komatiites are relatively common in the pre-Cambrian, however, the only known Phanerozoic komatiites occur as part of the CCOP on Gorgona Island. The formation of these Cretaceous komatiites in the CCOP has led to the suggestion that preCambrian komatiites formed in ancient oceanic plateaus (Storey et al., 1991). The lavas of the CCOP are classified as tholeiitic. The most magnesian lavas found in the province contain up to 28 wt.% MgO (Figure 5). However, as shown by Kerr et al.

(1996b), it is likely that these lavas contain substantial accumulated olivine, and so the whole rock compositions of these high-MgO rocks cannot represent those of parental mantle melts. Estimates of the MgO content of the parental melts for various parts of the province vary from 18 wt.% MgO to about 12 wt.% MgO (Kerr et al., 1996a,b; Revillon et al., 1999). Although picritic lavas are more common than in other Cretaceous oceanic plateaus, basalts are by far the most common rock type preserved in the CCOP. The vast majority of samples contain between 6 and 10 wt.% MgO (Figure 5). Al2O3 contents broadly increase with decreasing MgO, reflecting the importance of the addition and removal of olivine during the petrogenetic history of the CCOP lavas. CaO increases with decreasing MgO until MgO reaches 8 – 10 wt.% beyond which the CCOP lavas display a scattered but discernible downward trend. Fe2O3(t) and TiO2 display broadly horizontal trends until about 8 – 10 wt.% MgO, below which both increase markedly. These trends can be modeled by the initial fractional crystallization or accumulation of

546

Oceanic Plateaus

Figure 6 Plots of trace elements (ppm) and ratios of trace elements against wt.% MgO for lavas from the CCOP. The OJP is plotted as a single field except where the A and CG types differ markedly in composition. Data sources are as for Figure 5.

Geochemistry of Cretaceous Oceanic Plateaus

547

Figure 7 Primitive mantle normalized multi-element plot showing the average composition of high-MgO lavas (komatiites and picrites) and basalts from various parts of the CCOP, plotted along with average N-MORB values (dashed line) and a compositional field for the OJP. Data sources are as for Figure 5. Primitive mantle normalizing values and N-MORB from Sun and McDonough (1989).

olivine (plus minor Cr-spinel) followed by the commencement of crystallization of plagioclase and clinopyroxene between 8 wt.% MgO and 10 wt.% MgO (Kerr et al., 1996b). Trace element data (Figure 6) also support the proposed fractional crystallization model: Ni and Cr contents fall with decreasing MgO contents, and Sr, although commonly disturbed by subsolidus alteration processes, is generally reduced below 10 wt.% MgO. Despite some scatter, the content of incompatible trace elements, e.g., Nb and Zr, generally increases with decreasing MgO contents (Figure 6). Ratios of highly incompatible trace elements are not normally affected by moderate degrees of fractional crystallization or mantle melting, and they therefore have the potential to reveal heterogeneities in the mantle source region of the plateau. These ratios are plotted against MgO in Figure 6(g) – (h). One of the most interesting aspects of the trace element data for the CCOP is that the basaltic lavas possess a narrower range of incompatible trace element ratios than the picrites. For instance, well over 80% of the basaltic samples (,12 wt.% MgO) from the CCOP possess La/Y ratios between 0.05 and 0.2 and Zr/Nb ratios between 7 and 20 (Figure 6(g) –(h)). In contrast, the picritic and komatiitic lavas possess much more variable ratios of incompatible trace elements, with La/Y ranging from 0.05 to 0.45 and Zr/Nb from 5 to 85. This is also shown on primitive mantle normalized multielement plots, where it can be seen that the CCOP basalts possess broadly flat patterns whereas the high-MgO picrites and komatiites are generally much more variable, with some

being more depleted, and some more enriched than the basalts, particularly for the most highly incompatible trace elements such as Th, Nb, La, Ce & Nd (Figure 7). The heterogeneity of the high-MgO rocks is also reflected in the radiogenic isotope ratios, particularly 1Nd (Figures 8 and 9). Virtually all the analyzed basalts from the CCOP possess initial 1Nd ranging from þ6 to þ9, whereas the high MgO lavas generally fall outside this range (1Nd . þ9 and , þ6; Figure 8). Elevated initial 87 Sr/86Sr ratios found in several parts of the province have been attributed either to contamination with altered oceanic crust (Curac¸ao: Kerr et al., 1996b) or to secondary alteration (Gorgona: Revillon et al., 1999). The wide range of isotopic data for the CCOP reveals that the enriched and depleted lavas are not simply formed by variable melting of a homogeneous source region, but rather reflect melting of long-term depleted and enriched components from a markedly heterogeneous plume source region. (Kerr et al., 1996a, 2002; Arndt et al., 1997; Hauff et al., 2000a; Thompson et al., 2004). The greater heterogeneity of the high MgO rocks in comparison to the basalts, has been interpreted to reflect the formation of these lower MgO magmas through mixing and fractional crystallisation of the high-MgO magmas in large magma chambers. The heterogeneous high-MgO rocks thus represent magmas that passed relatively quickly through the lithosphere without being trapped in magma chambers (Kerr et al., 1998). The extent of partial melting required to produce

548

Oceanic Plateaus

Figure 8 Plot of (a) initial 1Nd against 87Sr/86Sr for high-MgO lavas and basalts from the CCOP and (b) 207Pb/204Pb against 206Pb/204Pb. Shown on both diagrams are fields for the OJP- A and CG types (Mahoney et al., 1993a,b) and East Pacific Rise (EPR) MORB (Mahoney et al., 1995). Other data sources are as for Figure 5. Mantle end-member compositions are from Zindler and Hart (1986).

the parental magmas of the CCOP has been calculated to be of the order of 20% (Kerr et al., 1997a). 3.16.4.4 Ontong Java Plateau (,122 and ,90 Ma) The Ontong Java plateau (OJP) in the western Pacific (Figure 1) has been tectonically uplifted and exposed along its south eastern margin, at the Solomon Islands arc, mostly on the Islands of Maliata and Santa Isabel. In contrast to the CCOP, which has numerous exposed sections, these are currently the only known subaerial exposures of the OJP. The rest of our knowledge of the OJP comes from a series of drill holes: DSDP Site 289 and ODP Sites 803 and 807 (Mahoney et al.,

Figure 9 Plots of (a) initial 1Hf against initial 1Nd and (b) initial g Os against initial 1Nd for Cretaceous oceanic plateaus. Data sources: Atlantic and Indian MORB—Salters (1996), Pacific—Nowell et al. (1998), Kerguelen plateau—Salters and Hart (1991), CCOP and Gorgona—Walker et al. (1999), Thompson et al. (2004), OJP—Babbs (1997).

1993a). Additional data are also now becoming available from ODP Leg 192 which recently penetrated the OJP at four sites (Mahoney et al., 2001). On Maliata the stratigraphic thickness of the accreted plateau reaches 3 –4 km, and the succession is dominated by pillowed and massive basaltic flows (Petterson et al., 1997; Babbs, 1997). Like the CCOP, dykes are volumetrically minor. The DSDP/ODP drill holes have penetrated into the plateau to a depth of 216 m (Site 1185B; Mahoney et al., 2001). The sampled sections consist predominantly of pillowed and massive basalts with occasional thin interlava sediments (Neal et al., 1997; Mahoney et al., 2001) In general, the OJP lavas are more homogeneous than those of the CCOP (Figures 5, 6 and 8). All the lavas from the province analyzed

The Influence of Continental Crust on Oceanic Plateaus thus far are basaltic in composition, with most of the samples possessing 6 –8 wt.% MgO (Figures 5 and 6). Although the lavas of the OJP possess a restricted compositional range, they nonetheless fall into two compositionally distinct groups. These groups were first noted in the lavas from ODP Site 807, where Mahoney et al. (1993b) divided the lavas into different units (A and C-G). Unit A (A-Type) is chemically distinct from Units C-G (C-G-Type) and possesses higher levels of both incompatible elements (e.g., TiO2; Sr, Zr Nb and the LREE; Figures 5 – 7) and ratios of highly incompatible to moderately incompatible trace elements (e.g., La/Y). The A-Type basalts also have lower initial 1Nd values and higher initial 87 Sr/86Sr ratios than the C-G-Type (Figure 8(a)). However, the total range of 1Nd (þ6.5 – þ4.9) and 87Sr/86Sr (0.7034– 0.7041) is relatively small in comparison with those in the CCOP (Figure 8(a)). The same is true for incompatible trace element ratios, with the OJP basalts only varying between 0.12 and 0.22 for La/Y and between 15.5– 19.7 for Zr/Nb (Figure 6). However, recent analyses from ODP Leg 192 (Sites 1185 and 1187; on the eastern edge of the plateau) have revealed the occurrence of more-MgO-rich lavas (up to 11 wt.%) with higher Ni and Cr contents (Mahoney et al., 2001; G. Fitton unpublished data; Figures 5 and 6). In addition to their higher MgO, preliminary geochemical data reveals that these basalts possess lower levels of incompatible elements (e.g., TiO2: 0.72 –0.77 and Zr: 36 –43), than the A and C-G Types (Figures 5 and 6). Neal et al. (1997) have concluded on the basis of geochemical modeling that the major and trace element compositions of the A- and C-G-type lavas of the OJP are consistent with 20– 30% partial melting of a peridotite source. The more enriched nature of the A-Type lavas implies derivation form a slightly more enriched source region, possibly in conjunction with smaller degrees of melting. Mahoney et al. (2001) have proposed that the more MgO-rich, incompatible element-poor lavas discovered during Leg 192 represent more extensive melting of the plume source region. However, an alternative explanation is that these lavas were derived from a more depleted mantle source region, and radiogenic isotope data are required in order to resolve this issue. None of the compositions sampled thus far are magnesian enough to represent possible parental melts, and so are believed to have undergone 30 – 45% fractional crystallization, involving olivine, plagioclase and clinopyroxene (Neal et al., 1997). Although the deeper crustal and lithospheric levels of the OJP are not exposed, seismic velocity data has been used to model the crustal structure (Farnetani et al., 1996; Gladczenko et al., 1997).

549

These authors have proposed that the magma chambers that fed the plateau are represented in the midcrust by olivine gabbros. These models also suggest that high compressional P-wave velocities of .7.1 km s21 deep within the OJP are due to the presence of olivine and pyroxene cumulates produced by the fractionation of primary picritic melts. Alternatively, the high P-wave velocities could be due to the presence of garnet granulite deep in the plateau, which Gladczenko et al. (1997) suggested may have formed by deformation and hydrothermal alteration of lower crustal cumulates.

3.16.5 THE INFLUENCE OF CONTINENTAL CRUST ON OCEANIC PLATEAUS Initially it may seem odd that the composition of oceanic plateaus should be influenced by continental crust, and certainly for the CCOP and the OJP, which apparently formed well away from continental margins, there is no evidence of the involvement of continental crust in their petrogenesis. However, LIPs can also form at the continent – ocean boundary as well as erupting onto either oceanic or continental lithosphere, and the formation of a LIP in such a tectonic setting is often related to continental break-up. The role played by mantle plumes in continental break-up (causal or consequential) remains controversial (White and McKenzie, 1989; Hill, 1991; Coffin and Eldholm, 1992; Saunders et al., 1992; Barton and White, 1995). However, whether mantle plumes are the reason for or a result of continental break-up, the associated erupted lavas and intruded sills form thick magmatic sequences on the margins of the rifted continents: the so-called seaward-dipping reflector sequences (SDRS). These LIPs may also erupt on the adjacent continents to form continental flood basalt provinces. Furthermore, continuing plume-related magmatism combined with further separation of the continents ultimately results in the formation of oceanic plateaus. Two examples of provinces such as these are explored below: the North Atlantic Igneous Province (NAIP) and the Kerguelen plateau. 3.16.5.1 The North Atlantic Igneous Province (,60 Ma to Present Day) The opening of the North Atlantic ,60 Ma is closely associated with magmatism from the “head” phase of the Icelandic plume. (For a comprehensive review of the NAIP see Saunders et al., 1997). Much of the initial volcanism (Phase 1: 62 –58 Ma; Saunders et al., 1997) was confined to the continental margins, i.e., the onland sequences in western Britain, the Faroe

550

Oceanic Plateaus

Islands and east and west Greenland, as well as the seaward-dipping reflector sequences of the southeast Greenland margin and the Hatton Bank (Figure 10). Most of these lavas are contaminated with Archean-age continental crust and thus possess low 1Nd and high Ba/Nb (Figure 11), along with low 206Pb/204Pb. As the North Atlantic continued to open, a second intense burst of magmatism occurred (beginning at 56 Ma; Phase 2, Saunders et al., 1997). The lavas from this magmatism are preserved in the upper portions of the SDRS, off the coast of southeast Greenland and Western Europe and have been drilled by ODP Legs 104, 152 and 163 (Viereck et al., 1988; Fitton et al., 2000). In contrast to the Phase 1 lavas, these lavas show few signs of contamination by continental crust (low Ba/Nb; 1Nd . 6; (Figure 11) 206Pb/204Pb .17), indicating that by this time the NAIP was an entirely oceanic LIP. The Icelandic plume has been producing melt over most of the past 60 Myr, as evidenced by 55 – 15 Ma volcanism along the Greenland –Iceland ridge and the Faroes –Iceland ridge, and the 15 Ma – present volcanism on Iceland.

3.16.5.2 The Kerguelen Igneous Province (,133 Ma to Present Day) The initial volcanism of the Kerguelen plume is closely associated with the break-up of Gondwana in the early – mid-Cretaceous, i.e., the separation of India, Australia and Antarctica (Morgan, 1981; Royer and Coffin, 1992). Like the NAIP, much initial volcanism is found on the margins of the rifted continents (Figure 12): the Rajmahal basalts in northeastern India (Kent et al., 1997) and the Bunbury basalts in western Australia (Frey et al., 1996). Not surprisingly, these basalts are extensively contaminated by continental lithosphere and yield an initial 87 Sr/ 86 Sr ratio of .0.7042 and 1Nd , 4.0 (Figure 13). The geographical components of the plateau (Figure 12) and the geochronology are briefly outlined below; however, a more detailed review can be found in Frey et al. (2000); Coffin et al. (2002). The first massive pulse of Kerguelen plume magmatism created the Southern Kerguelen plateau (118 – 110 Ma; Figure 12). Later melting of the plume was responsible for

Figure 10 Map showing the locations of the principal on-land exposures of the North Atlantic Igneous Province and the seaward-dipping reflector sequences.

Identification of Oceanic Plateaus in the Geological Record

551

Figure 11 Plot of Ba/Nb vs. initial 1Nd lavas from the NAIP. Shown on the diagram are lavas from Skye (Thompson et al., 1982; Dickin et al., 1987), Mull (Kerr et al., 1995), Iceland (He´mond et al., 1993) and both pre- and post- continental break-up lavas from ODP Leg 152 (Fitton et al., 1998a).

the formation of the Elan Bank (108 – 107 Ma), the Central Kerguelen Plateau (101 – 100 Ma), Broken Ridge (95 – 94 Ma), the Ninetyeast Ridge (82 –37 Ma), and the Northern Kerguelen plateau (35 –34 Ma) (Figure 12). Volcanism continues to the present-day and has produced the Kerguelen Archipelago and Heard and MacDonald Islands. The lavas from the Southern Kerguelen plateau and Broken Ridge have initial 87Sr/86Sr ratios and 1Nd values (Figure 11) which range from 0.7037 to 0.7102 and þ4.0 to 29.4, respectively (Salters et al., 1992; Mahoney et al., 1995). Some of this variation can be interpreted as mixing between Southeast Indian Ridge MORB and the Kerguelen plume (Weis and Frey, 1996). However, elevated La/Nb ratios (Figures 13 and 14(a)) and the extreme isotopic compositions of basalts drilled at ODP Site 738 and dredge samples from the eastern Broken Ridge ( 87 Sr/ 86Sr 0.710; 1Nd 2 9.0; Figures 8 and 13) cannot be explained by such mixing processes. It has been proposed that these signatures are due to contamination by continental lithosphere (Storey et al., 1989; Mahoney et al., 1995; Hassler and Shimizu, 1998). Operto and Charvis, 1996 have imaged a seismically reflective transition zone beneath the crust/mantle interface of the Southern Kerguelen plateau, interpreted as fragments of continental crust. This crust appears to have isotopic similarities to Archean crust found on the margins of Gondwana, which raises the possibility that fragments of such crust have become incorporated into the Indian Ocean basin during continental break-up. Recently, during drilling at Site 1137 on the Elan Bank (part of the Kerguelen plateau; Figure 12) clasts of garnet-biotite gneiss have been discovered in a fluvial conglomerate

Figure 12 Map showing the main components of the Kerguelen plateau (KP) discussed in the text (after Frey et al., 1996).

intercalated with basalt flows (Frey et al., 2000). This discovery has confirmed the presence of preCambrian crustal rocks within the Kerguelen plateau (Nicolaysen et al., 2001), thus supporting the lithospheric contamination model for the high La/Nb, low 1Nd basalts of the Kerguelen plateau.

3.16.6 IDENTIFICATION OF OCEANIC PLATEAUS IN THE GEOLOGICAL RECORD The rationale for this section is summed up by this question: If the CCOP or OJP were accreted on to a continental margin and preserved in the geologic record for 1 billion years, what features could we use to identify them as oceanic plateaus? This section will review diagnostic geochemical and geological characteristics of oceanic plateaus, and then will show, illustrated by examples, how these criteria can be used to identify plateau sequences in the geological record. Table 3 provides a summary of the diagnostic features of Cretaceous oceanic plateaus and mafic sequences

552

Oceanic Plateaus

Figure 13 Plots to show the geochemical variation of lavas from the early Cretaceous lavas derived from the Kerguelen plume. (a) Initial 87Sr/86Sr vs. initial 1Nd and (b) primitive mantle normalized multi-element plots showing averaged data for ODP drill sites. A compositional field for the OJP (Mahoney et al., 1993a,b) is shown on both diagrams. Data sources: Rajmahal—Kent et al. (1997); Bunbury—Frey et al. (1996); ODP sites—Salters et al. (1992); Mahoney et al. (1995).

within the continental crust, which have been interpreted as oceanic plateaus. Details of accreted oceanic plateaus thus far identified in the geological record are summarized in Table 4. 3.16.6.1 Diagnostic Features of Oceanic Plateaus Both chemical and geological features can be useful in the identification of oceanic plateaus. Condie (1999) and Kerr et al. (2000) have discussed the diagnostic features of oceanic plateaus in detail, and only a brief account will be given here. Table 3 summarizes the characteristics which are useful in distinguishing igneous rocks formed in an oceanic plateau from those which originated in other tectonic settings. Igneous rocks produced in an island arc, or a continental subduction zone setting, are relatively easily distinguished from oceanic plateau sequences (Table 3), because arcs generally possess more evolved lavas, with ubiquitous high (La/Nb)pmn ratios (Figure 14(b)), and only very rarely contain high-MgO lavas. Additionally,

oceanic plateaus do not possess the abundant volcanic ash layers present in volcanic arc sequences. However, as Figure 14 shows, a low (La/Nb)pmn ratio is not an entirely robust signature of an oceanic plateau sequence, since samples of the Kerguelen oceanic plateau often possess high (La/Nb)n values, due to magma interaction with fragments of continental lithosphere beneath the plateau. This example highlights the importance of not relying solely on chemical discriminants of tectonic environment, without also considering the geological evidence. In the case of Kerguelen, the lack of volcaniclastic horizons helps confirm its oceanic plateau affinity. As discussed by Kerr et al. (2000), many of the geological discriminants between oceanic plateaus and midocean ridges may be ambiguous (Table 3). Geochemical characteristics must, therefore, be used to distinguish lavas from these two tectonic settings. Most Cretaceous oceanic plateau lavas possess relatively flat normalized REE patterns (Figure 7), whereas most midocean ridge basalts possess light REE-depleted patterns reflecting a more depleted

Identification of Oceanic Plateaus in the Geological Record

553

record simply because they are easily eroded away, unless they are buried by sediments. Quite often the only remaining indications of a continental flood basalt province are the dykes and vents through which the lavas erupted. The 200 Ma Central Atlantic Magmatic Province, which formed during the break-up of South America, Africa, and North America has been identified largely on the basis of its remnant dyke swarms (e.g., Marzoli et al., 1999).

3.16.6.2 Mafic Triassic Accreted Terranes in the North American Cordillera

Figure 14 (a) Frequency diagram showing the range in (La/Nb)pmn for lavas from the CCOP– OJP, the Kerguelen plateau, EPR MORB, and arc lavas. (Arc data from Thirlwall et al., 1996.) Other data sources as in Figures 5, 8 and 12). (b) Plot of Ni vs. Mg – number for the CCOP – OJP and lavas from back arc basins. Data from Wood et al. (1980), Woodhead et al. (1998), Leat et al. (2000). Both diagrams modified from Kerr et al. (2000).

mantle source region. Furthermore, high-MgO lavas can be found in oceanic plateaus, but are largely absent from oceanic crust generated at midocean ridges. Incompatible trace elements are only of limited use in distinguishing between volcanic successions formed in back-arc basins and those formed in oceanic plateaus (Table 3). However, the lower mantle temperature below a back-arc basin (Tp , 1,280 8C) compared to a mantle plume (Tp . 1,400 8C) results in the eruption of few high-MgO lavas. An additional consequence of this lower mantle temperature is that back-arc basin lavas generally possess lower Ni and Cr contents for a given Mg number than oceanic plateau lavas (Figure 14(b)). Furthermore, because of their proximity to active subduction sites, back-arc basin sequences are also more likely to contain abundant volcaniclastic horizons than oceanic plateaus. Continental flood basalts are not easy to preserve for long periods of time in the geological

Significant proportions of the North American Cordillera consist of mafic sequences of accreted oceanic terranes (Figure 15). Some of these have been identified as oceanic plateau material ranging in age from Permian to Eocene (see review in Condie, 2001). At least three of these oceanic plateau terranes are predominantly Triassic in age (Wrangellia, Cache Creek and Angayucham; Pallister et al., 1989; Lassiter et al., 1995; Tardy et al., 2001) and obviously represent a major phase of oceanic plateau volcanism at this time. These plateau sequences are characterized by pillow basalts and intrusive sheets, with occasional intercalated tephra and hyaloclastite layers, indicating formation in shallow water, or by subaerial eruption. In the Wrangellia terrane there is considerable evidence for rapid uplift of the sea floor (presumably by the plume head) immediately prior to eruption (Richards et al., 1991). The basalts of the Cache Creek and Angayucham terranes display a restricted range in MgO with most of the basalts ranging from 5.0 wt.% to 8.5 wt.%. These basalts possess low (La/Nb)pmn ratios (,1.2), essentially flat REE patterns (Pallister et al., 1989; Tardy et al., 2001) and 1Nd values that range mostly from þ9.9 to þ4.5 (Figure 16). As Figure 16 shows, all these features are similar to the OJP. However, some of the basalts from the Wrangellia Terrane, despite showing a similar range in MgO content, have (La/Nb)pmn ratios .1 and steeper REE patterns (Figure 16) than those from Cache Creek and Angayucham. Lassiter et al. (1995) suggested that this is due to the magmas erupting through, and being contaminated by preexisting island-arc lithosphere. However, it is also possible that, as in the case of the Kerguelen plateau, large fragments of ancient continental lithosphere were incorporated in the proto-Pacific Ocean, and the lavas of the Wrangellia oceanic plateau were contaminated by this lithosphere. The contamination of the most evolved Wrangellia basalts with either arc crust or ancient continental lithosphere is also supported by a broadly negative correlation between 1Nd and (La/Nb)pmn (Figure 16).

Oceanic Plateaus

no yes occasional

Rare frequent yes

yes no no rare no occasional rare no frequent frequent rare rare yes rare occasional

Flat to LREE-enriched mostly . . 1 , 10% of flows #1 a

pmn—primitive mantle normalized (after Kerr et al. 2000).

frequent frequent Continental flood basalt

Arc (continental & oceanic)

rare

frequent

LREE-enriched

Predominantly flat LREE-depleted Predominantly flat Predominantly LREE-enriched Flat to LREE-enriched Oceanic plateau Midocean ridge Marginal basin Oceanic island basalt Volcanic rifted margin

frequent rare rare rare frequent

rare rare rare rare rare

#1 #1 #1 #1 Varies from #1 to . .1 . .1

yes yes yes yes not all lavas are pillowed not all lavas are pillowed no

Subaerial eruption Tephra layers Chondrite normalized REE pattern

Pillow lavas

3.16.6.3 Carboniferous to Cretaceous Accreted Oceanic Plateaus in Japan

(La/Nb)apmn Low-MgO lavas (,3%) High-MgO lavas (.14%) Tectonic setting

Table 3 Diagnostic geochemical and geological characteristics of volcanic sequences from different tectonic settings.

Intercalated pelagic sediments

554

The Japanese islands are essentially composed of a series of terranes that have been accreted to the continental margin of the Eurasian plate during the past 400 Myr. These terranes consist of trench-filling terrigenous sediments with variable quantities of accreted oceanic crust that are intruded and partly overlain by the products of subsequent subduction-related volcanism. Within Japan the ages of the accreted complexes become younger from north to south and from west to east (Kimura et al., 1994). However, relatively little is known about the trace element chemistry of these oceanic accreted terranes. The Chugoku and Chichibu belts in southwest Japan contain up to 30% basaltic material (greenstones) in thrust contact with limestones, cherts and mudstones (Tatsumi et al., 2000). This lithological association, combined with preliminary major element data and a small range of trace element data, suggests that these basaltic assemblages are remnants of a plume-derived oceanic plateau. Tatsumi et al. (2000) proposed that this oceanic plateau formed in the Panthalassan Ocean (proto-Pacific) in the Carboniferous (350 –300 Ma). The Permian Yakuno ophiolite complex in southern Honshu Island is a sequence of submarine basalts, gabbros and ultramafic rocks (Isozaki, 1997). The presence of pelagic sediments, the lack of a sheeted dyke complex and the fact that the sequence is of considerable thickness, all suggest that it is part of an oceanic plateau (Isozaki, 1997). In contrast to the other accretionary belts in Japan, the Sorachi-Anivia terrane in Hokkaido and Sakhalin is dominated by oceanic crust and lithosphere (Kimura et al., 1994; Tatsumi et al., 1998). It comprises pillow lavas and dolerite sills with intercalated pelagic sediments containing Tithonian (150 – 145 Ma) radiolaria, along with a lower unit in which ultramafic rocks, including serpentinite, harzburgite and dunite, are more common (Kimura et al., 1994). Major element data have shown that some picrites (.12 wt.% MgO) are found within the succession. This, combined with high CaO/Al2O3 ratios (indicative of a high degree of mantle melting), led Kimura et al. (1994) and Tatsumi et al. (1998) to propose an oceanic plateau origin for the Sorachi-Anivia terrane. This Jurassic – early Cretaceous plateau (named the Sorachi plateau; Kimura et al., 1994) is of the same age as the Shatsky Rise (Figure 1, Table 1). In combination with paleomagnetic data, this suggests that the Sorachi plateau and the Shatsky Rise were originally a single plateau which formed near the Kula – Pacific – Farallon triple junction ,150 Ma (Kimura et al., 1994).

Table 4 Proposed accreted oceanic plateaus found within continents. Name

Location

Age (Ga)

References

Coonterunah and Warrawoona Groups Southern Barberton Belt Pietersberg Belt Opapimiskan-Markop Unit, North Caribou Belta Olondo Belt

Pilbara Craton, Australia Kaapvaal Craton, S Africa Kaapvaal Craton, S Africa Superior Province Aldan Shield, Siberia

,3.5 3.5– 3.2 ,3.4 ,3.0 3.0

South Rim Unit, North Caribou Belt Sumozero-Kenozero Belt Steep Rock & Lumby Lake Beltsa Balmer Assemblage, Red Lake Greenstone Belta Kostomuksha Belt Vizien Belt Malartic-Val d’Or Area Tisdale Group, Abitibi Belt Schreiber-Hemlo-White River Dayohessarah Vetreny Belta Birimian Province Povungnituk & Chukotat Groupsa

Superior Province Baltic Shield Superior Province Superior Province

,3.0 3.0– 2.8 3.0– 2.9 2.99– 2.96

Green et al. (2000) De Wit et al. (1987) De Wit et al. (1987) Hollings and Wyman (1999) Puchtel and Zhuravlev (1993); Bruguier (1996) Hollings and Wyman (1999) Puchtel et al. (1999) Tomlinson et al. (1999) Tomlinson et al. (1998)

Baltic Shield Superior Province Superior Province Superior Province Superior Province Baltic Shield West Africa Cape Smith Fold Belt Northern Que´bec Baltic Shield NE Finland Central Canada NE Africa-Middle East SW Japan SW Japan Central Japan Canadian Cordillera Alaska Western North America Northern Japan

2.8 2.79 2.7 ,2.7 2.8– 2.7 2.44 2.2 2.04

Puchtel et al. (1998b) Skulski and Percival (1996) Kimura et al. (1993); Desrochers et al. (1993) Fan and Kerrich (1997) Polat et al. (1998) Puchtel et al. (1997) Abouchami et al. (1990); Boher et al. (1992) Francis et al. (1983); Dunphy et al. (1995)

1.98 1.95 1.92– 1.90 0.90-0.87 Carboniferous 0.285 L Permian Triassic Triassic 0.227 0.152– 0.145

Puchtel et al. (1998a) Peltonen et al. (1996) Lucas et al. (1996); Stern et al. (1995) Stein and Goldstein (1996) Tatsumi et al. (2000) Isozaki (1997) Jones et al. (1993) Tardy et al. (2001) Pallister et al. (1989) Lassiter et al. (1995). Kimura et al. (1994); Tatsumi et al. (1998)

Onega Plateaua Jormua Ophiolitea Flin Flon Belt Arabian-Nubian Shield Chichibu & Chugoku Belts Yakuno Ophiolite Mino Terrane Cache Creek Terrane Angayucham Terrane Wrangellia Terrane Sorachi Plateau a

These sequences display evidence of contamination by continental crust and are interpreted as having formed during continental break-up or, close to a continental margin (see text).

556

Oceanic Plateaus The Sorachi part of the plateau was carried northwestwards on the Kula plate and collided with Japan ,110 Ma. Limited trace element data for the Sorachi plateau lavas support a common plume source for these two plateaus. The data cover the same compositional range as that of dredged samples from the Shatsky Rise (Figure 17). Furthermore, the data reveal that the plume source region of the Sorachi plateau was markedly heterogeneous and contained both enriched and depleted components (Kimura et al., 1994; Tatsumi et al., 1998) (Figure 17).

3.16.7 PRECAMBRIAN OCEANIC PLATEAUS Figure 15 Map showing North American accreted oceanic terranes including the oceanic plateau sequences discussed in the text (after Tardy et al., 2001; Condie, 2001).

The identification of accreted pre-Cambrian oceanic plateaus, particularly in greenstone belts, has important implications for the generation of continental crust (Abbott, 1996; Albarede, 1998; Condie, 1999). Kerr et al. (2000) have presented

Figure 16 Plots of (a) (La/Nb)pmn against initial 1Nd for Cache Creek, Wrangellia, OJP and the Kerguelen plateau and (b) chondrite normalized (Sun and McDonough, 1989) REE plot showing averages for Wrangellia, Angayucham, Cache Creek, and the Kerguelen plateau, with the range for the OJP. Data sources are as in Figures 5 and 12; North American data from Pallister et al. (1989); Lassiter et al. (1995); and Tardy et al. (2001).

Environmental Impact of Oceanic Plateau Formation

Figure 17 Plot of Nb/Y against Nb/Zr showing data from dredge samples from the Shatsky Rise and the accreted Sorachi plateau (data from Tatsumi et al., 1998). Also shown are fields for the CCOP and OJP. Data sources are as in Figure 5.

a summary of accreted pre-Cambrian oceanic plateaus and the reader is referred to their paper for more detailed information. Some of the oldest preserved oceanic plateau sequences are those found in ,3.5 Ga Barberton and Pietersberg belts of the Kaapvaal Shield of southern Africa (De Wit et al., 1987; Smith and Erlank, 1982). These belts contain pillow basalts and komatiites, with chemical signatures (Lahaye et al., 1995) suggesting a likely origin as part of an oceanic plateau. The Pilbara craton of Australia appears to possess some of the oldest oceanic plateau material so far identified (Green et al., 2000) in the ,3.5 Ga Coonterunah and Warrawoona Groups. Greenstone belts of the Canadian Superior province, ranging in age from 3.0 to 2.7 Ga also contain lava groups that have been interpreted to be remnants of accreted oceanic plateaus. These belts include the South Rim unit of the North Caribou belt (Hollings and Wyman, 1999), the Vizien belt (Skulski and Percival, 1996), the Malartic-Val d’Or (Desrochers et al., 1993), the Tisdale Group of the Abitibi belt (Fan and Kerrich, 1997), and the Schreiber-Hemlo and the White River-Dayohessarah belts (Polat et al., 1998). The evidence for an oceanic plateau origin is based on the occurrence of pillow basalts and komatiites without terrestrial sedimentary intercalations or sheeted dyke swarms, possessing low (La/Nb)pmn and the low positive 1Nd (Figure 18) that are characteristic of Cretaceous oceanic plateaus. Several of these sequences within the Superior province are in stratigraphic contact with basalts and komatiites that possess a signature of continental lithosphere contamination, i.e., negative 1Nd and (La/Nb)pmn .1 (Figure 18). These

557

Figure 18 Plot of (La/Nb)pmn vs. initial 1Nd for data from Archean and Proterozoic accreted oceanic plateaus and provinces proposed to have formed during continental rifting. Data sources: Abouchami et al. (1990); Skulski and Percival (1996); Puchtel et al. (1998a); Tomlinson et al. (1998); Hollings and Wyman (1999).

include the Balmer assemblage, Red Lake belt (Tomlinson et al., 1998), Steep Rock-Lumby Lake belts (Tomlinson et al., 1999), and OpapimiskanMarkop unit, North Caribou belt (Hollings and Wyman, 1999). These sequences have been interpreted as having formed in tectonic settings related to continental break-up, similar to the North Atlantic Tertiary igneous province and parts of the Cretaceous Kerguelen plateau. Puchtel et al. (1998b, 1999) have proposed that the 3.0 – 2.8 Ga Kostomuksha and SumozeroKenozero greenstone belts of the Baltic Shield represent remnants of oceanic plateaus. This interpretation is based on the occurrence of crustally uncontaminated pillow basalts (Figure 18) and komatiites without terrestrial sedimentary intercalations. In contrast the 2.4 Ga Vetreny greenstone belt and the 1.98 Ga Onega plateau, also part of the Baltic Shield, display chemical evidence of crustal contamination (negative 1Nd; Figure 18). These sequences were interpreted by Puchtel et al. (1997, 1998a) as having formed during continental break-up. Other Proterozoic oceanic plateau terranes have been identified in the Birimian province of western Africa (Figure 18) (Abouchami et al. 1990; Boher et al., 1992), the Arabian-Nubian Shield (Stein and Goldstein, 1996), and the Flin Flon belt in Canada (Stern et al., 1995). 3.16.8 ENVIRONMENTAL IMPACT OF OCEANIC PLATEAU FORMATION Although the potential environmental impact of continental flood basalt provinces has been

558

Oceanic Plateaus

documented by many authors (e.g., Hallam, 1987a; McLean, 1985; Renne and Basu, 1991; Courtillot et al., 1996), the possible effects of oceanic plateau eruptions on the atmosphere, biosphere and hydrosphere have received comparatively little attention (see, for example, Courtillot, 1994). This omission is surprising since the inclusion of oceanic plateau events actually strengthens the correlation between LIP eruptions and mass extinction events (Kerr, 1997), by providing a feasible terrestrial causal mechanism for several second order extinction events (Sepkoski, 1986). 3.16.8.1 Cenomanian–Turonian Boundary (CTB) Extinction Event Several of these second order extinction events occurred in the mid-Cretaceous. One of these, the CTB event (,93 Ma) has been linked by several authors to the formation of oceanic plateaus (Sinton and Duncan, 1997; Kerr, 1998). The CTB event is characterized by the world-wide deposition of organic-rich black shales (Jenkyns, 1980; Schlanger et al., 1987). The formation of black shale implies a widespread reducing environment (“anoxia”) in the oceans at this time (Figure 19). In addition to this, the CTB was a time of major sea level transgression (Hallam, 1989) and is marked by a positive carbon isotopic anomaly (d 13C excursion) of up to þ4 –5‰ (Arthur et al., 1987), indicating an increase in organic carbon burial rate (Figure 19). Sea water 87Sr/86Sr (Figure 19) reaches a maximum of 0.70753 in the late-Cenomanian, and drops steadily to a value of 0.70737 in the midTuronian, before starting to rise again. Average global surface temperatures (including oceanic temperatures) around the CTB were 6 – 14 8C higher than present (Kaiho, 1994), and this is most likely due to an increase in global atmospheric CO2 content which may have been .10 times present-day levels (Figure 19) (Arthur et al., 1987). These phenomena were accompanied by an extinction event that resulted in the demise of 26% of all known genera (Sepkoski, 1986). Although the overall extinction rate is much lower than that at the Cretaceous– Tertiary boundary, deep water marine invertebrates fared much worse in the CTB event (Kaiho, 1994). This difference supports the view that anomalous oceanic volcanism around the CTB may have played a significant role in the environmental and biotic crisis at this time (Kerr, 1998). Siderophile and compatible lithophile trace elements such as Sc, Ti, V, Cr, Mn, Co, Ni, Pt, Ir and Au are enriched in CTB black shales (Leary and Rampino, 1990; Orth et al., 1993). Kerr (1998) has shown that trace element abundances

Figure 19 Graphs showing how various parameters discussed in the text vary from 110 Ma to 80 Ma. The dotted vertical line represents the Cenomanian – Turonian boundary (after Kerr, 1998).

and ratios found in CTB black shales are similar to plume-derived volcanic rocks and midocean ridge basalts. For example, in mafic volcanic rocks Ni/Ir £ 104 values range from 70 to 190, and in CTB sediments this ratio averages 180. In contrast, average sedimentary rocks possess Ni/Ir £ 104 ratios of ,100 (Orth et al., 1993).

3.16.8.2 Links between CTB Oceanic Plateau Volcanism and Environmental Perturbation The most extensive plume-related volcanism around the CTB occurred in the oceans, with

Environmental Impact of Oceanic Plateau Formation the formation of the CCOP along with portions of the OJP and Kerguelen plateau. In addition to this oceanic volcanism, a continental flood basalt province related to the Marion hotspot also erupted at this time, as Madagascar rifted from India (Storey et al., 1995). The estimated erupted volume of oceanic plateau lavas around this time is ,1.0 £ 107 km3 and may be much higher (Kerr, 1998). The potential physical and chemical effects of oceanic plateau volcanism on the global environment are summarized in Figure 20 and discussed below. An obvious physical effect of oceanic mantle plume volcanism is to raise sea level by lava extrusion onto the ocean floor through the buoyant plume head uplifting the oceanic lithosphere and

Figure 20

559

displacing seawater (Courtney and White, 1986) and by the thermal expansion of seawater due to heating. The steady rise in global sea level throughout the Late Albian and Cenomanian (Figure 19) may reflect the arrival of the Caribbean, Ontong Java and Kerguelen plume heads below the oceanic lithosphere, prior to extensive volcanism (Vogt, 1989; Larson, 1991). This plume-related uplift of oceanic lithosphere may also have caused the disruption of important oceanic circulation systems such that cool, polar (oxygenated) waters were not circulated to lower latitudes, resulting in increased oceanic anoxia. Additionally, hydrothermal fluids from oceanic plateau volcanism could have contributed to warmer oceans, and thus to anoxia, since oxygen

Flow diagram summarizing the possible physical and chemical environmental effects of the formation of large igneous provinces around the Cenomanian –Turonian boundary (after Kerr, 1998).

560

Oceanic Plateaus

solubility is considerably reduced in warmer seawater (Sinton and Duncan, 1997). Positive d 13C anomalies at the CTB reflect increased rates of organic carbon burial as a result of high productivity and more effective preservation of organic material (Arthur et al., 1987). Such increased productivity means the supply of deep ocean nutrients, such as phosphates, into surface waters must also increase and this process may have been induced by oceanic plateau volcanism (Vogt, 1989; Sinton and Duncan, 1997). Elevated CO2 levels at the CTB may also be due to increased volcanic activity. Kerr (1998) has calculated that approximately 1017 kg of CO2 would have been released as a result of oceanic plateau volcanism around the CTB (Arthur et al., 1987). Additionally, LIP volcanism also releases substantial amounts of SO2, chlorine, fluorine and H2S which, when released into seawater, would have made the CTB oceans much more acidic (Kerr, 1998). The lack of carbonate at the CTB may be the result of increased dissolution by more acidic seawater, which would also release more CO2 to the atmosphere (Arthur et al., 1987). These additions of CO2 to the atmosphere would have resulted in significant global warming. The solubility of CO2 in seawater decreases as the temperatures rise; so the warmer the oceans get, the less CO2 will dissolve in them. Thus, with this positive CO2 feedback mechanism it is possible that a “runaway greenhouse” climate may have developed quite rapidly (Figure 20) (Kerr, 1998). Increased weathering of continental silicates can reduce CO2 levels. However, the rate of CO2 release at the CTB was much greater than its uptake by slow weathering processes. Increased atmospheric CO2 and the upwelling of nutrients from the deep ocean could have resulted in increased productivity in ocean surface waters (Figure 20), leading to the widespread deposition of black shales and thus a reduction in CO2 levels. Increased concentration of toxic trace metals in the oceans, liberated by hydrothermal fluids from oceanic plateau lava piles, may well have been a contributory factor to the demise of some marine organisms around the CTB (Wilde et al., 1990). The upwelling of deep ocean trace metals and

nutrients may result in the enlargement of the trace metal-restricted habitat of deeper dwelling organisms (Wilde et al., 1990), leading to increased predation by deeper dwelling creatures on those living in shallower water. Throughout the past 250 Myr significant black shale deposits occur during several other periods and, like the CTB, these other black shales are associated with, sometimes severe, environmental disruption (Jenkyns, 1980; Hallam, 1987b; Arthur and Sageman, 1994). It is interesting, and probably highly significant, that these black shale events correlate with the formation of oceanic plateaus or plume-related volcanic rifted margins (Table 5). The Aptian– Albian (121 – 99 Ma) appears to have been a period of persistent environmental disturbance with three distinct oceanic anoxia events (with associated black shales) during this period (Bralower et al., 1993). A causal link between black shale formation, environmental perturbation and oceanic volcanism is given further credence by the fact that a major period of oceanic plateau formation occurred during the Aptian –Albian (121 – 98 Ma) (see above). Finally, Condie et al. (2001) have presented evidence that the correlation between black shale deposition, paleoclimatic disturbance and mantle superplume events can be extended back to the pre-Cambrian. Particularly good correlations between environmental disturbance and mantle plume activity occur at 1.9 Ga and 2.7 Ga

3.16.9 CONCLUDING STATEMENTS Oceanic plateaus represent overthickened areas of oceanic crust (10 – 35 km), which appear to have formed as a result of decompression melting of a large mantle plume head, often (although not always) within 1 –2 Ma. Geological and geochemical evidence suggests that oceanic plateaus have formed throughout a considerable period of Earth’s history. The thickness of the crustal sections of oceanic plateaus implies that they are not easily subducted. Thus, when these plateaus encounter a subduction zone, their top-most portions tend to become

Table 5 Correlations between black shale events and oceanic plateau volcanism over the last 250 Ma. Age

Black shales

Aptian– Albian (121 – 99 Ma)

Extensive world-wide deposits

Tithonian (150 – 144 Ma)

Extensive deposits in Europe and west Asia Extensive world-wide deposits Few deposits

Toarcian (190 – 180 Ma) Carnian (227– 220 Ma)

Oceanic plateau or volcanic rifted margin Ontong Java & other Pacific plateaus, Kerguelen Sorachi plateau & Shatsky Rise Karoo, Ferrar & Weddell Sea Wrangellia

Data sources: Hallam (1987b); Jenkyns (1980); Riley and Knight (2001); Hergt and Brauns (2001).

References accreted to a continent or an island arc. In this way fragments of oceanic plateaus have become incorporated into continental margins and preserved in the geological record. Two recent examples of such accreted oceanic plateaus are the CCOP and the OJP. All of the lavas sampled from the OJP and most of the lavas from the CCOP are relatively homogeneous basalts with initial 1Nd values between þ5 and þ8 and broadly chondritic trace element ratios. In contrast, high-MgO lavas found in the CCOP reveal evidence of a more heterogeneous plume source region, containing both enriched (1Nd , þ5; (La/Nd) cn .1) and depleted (1Nd , þ8; (La/Nd)cn ,1) components. The formation of LIPs during continental break-up (e.g., the NAIP and the Kerguelen plateau) results in the formation of seawarddipping reflector sequences and oceanic plateaus, both of which show chemical evidence of contamination by ancient continental lithosphere (initial 1Nd , 0; (La/Nb)pmn .1). The correlation between oceanic plateau formation and marine environmental catastrophes (characterized by mass extinction, oceanic anoxia and black shale deposition) throughout the Mesozoic period suggests a causal link between oceanic plateau formation and environmental crises.

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Thompson R. N., Dickin A. P., Gibson I. L., and Morrison M. A. (1982) Elemental fingerprints of isotopic contamination of Hebridean Palaeocene mantle-derived magmas by Archaean sial. Contrib. Mineral. Petrol. 79, 159–168. Thompson P. M. E., Kempton P. D., White R. V., Kerr A. C., Tarney J., Saunders A. D., and Fitton J. G. (2004) Hf-Nd isotope constraints on the origin of the Cretaceous Caribbean plateau and its relationship to the Galapagos plume. Earth Planet. Sci. Lett. (in press). Tomlinson K. Y., Stevenson R. K., Hughes D. J., Hall R. P., Thurston P. C., and Henry P. (1998) The Red Lake greenstone belt, Superior Province: evidence of plumerelated magmatism at 3 Ga and evidence of an older enriched source. Precamb. Res. 89, 59–76. Tomlinson K. Y., Hughes D. J., Thurston P. C., and Hall R. P. (1999) Plume magmatism and crustal growth at 2.9 to 3.0 Ga in the Steep Rock and Lumby Lake area, Western Superior Province. Lithos 46, 103– 136. Viereck L. G., Taylor P. N., Parson L. M., Morton A. C., Hertogen J., Gibson I. L., and Party O. S. (1988) Origin of the Palaeogene Vøring Plateau volcanic sequence. In Early Tertiary Volcanism and the Opening of the NE Atlantic, Geological Society London Special Publication 39 (eds. A. C. Morton and L. M. Parson), pp. 69– 83. Vogt P. R. (1989) Volcaniogenic upwelling of anoxic, nutrientrich water: a possible factor in carbonate-bank/reef demise and benthic faunal extinctions. Bull. Geol. Soc. Amer. 101, 1225–1245. Walker R. J., Storey M. J., Kerr A. C., Tarney J., and Arndt N. T. (1999) Implications of 1870s isotopic heterogeneities in a mantle plume: evidence from Gorgona Island and Curac¸ao. Geochim. Cosmochim. Acta 63, 713– 728. Weis D. and Frey F. A. (1996) Role of the Kerguelen Plume in generating the eastern Indian Ocean seafloor. J. Geophy. Res. 101, 13831–13849. White R. S. and McKenzie D. P. (1989) Magmatism at rift zones: the generation of volcanic continental margins. J. Geophy. Res. 94, 7685–7729. White R. V., Tarney J., Kerr A. C., Saunders A. D., Kempton P. D., Pringle M. S., and Klaver G. T. (1999) Modification of an oceanic plateau, Aruba, Dutch Caribbean: implications for the generation of continental crust. Lithos 46, 43–68. Wilde P., Quinby-Hunt M. S., and Berry B. N. (1990) Vertical advection from oxic or anoxic water from the pycnocline as a cause of rapid extinction or rapid radiations. In Extinction Events in Earth History. Lecture Notes in Earth Science (eds. E. G. Kauffman and O. H. Walliser). Springer, vol. 30, pp. 85 –97. Wilson J. T. (1963) A possible origin of the Hawaiian Islands. Can. J. Phy. 41, 863–870. Wood D. A., Mattey D. P., Joron J. L., Marsh N. G., Tarney J., and Treuil M. (1980) A geochemical study of 17 selected samples from basement cores recovered at Sites 447, 448, 449, 450, and 453, Deep Sea Drilling Project Leg 59. In Initial Reports of the Deep Sea Drilling Project 59 (eds. L. Kroenke, R. Scott, and et al.) U.S. Government Printing Office, pp. 743–752. Woodhead J. D., Eggins S. M., and Johnson R. W. (1998) Magma genesis in the New Britain island arc: further insights into melting and mass transfer processes. J. Petrol. 39, 1641–1668. Zindler A. and Hart S. R. (1986) Chemical geodynamics. Ann. Rev. Earth Planet. Sci. 14, 493–571.

q 2003, Elsevier Ltd. All rights reserved No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without prior written permission of the Publisher.

Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 537–565

3.17 Generation of Mobile Components during Subduction of Oceanic Crust M. W. Schmidt ETH Zu¨rich, Switzerland and S. Poli Universita di Milano, Italy 3.17.1 INTRODUCTION 3.17.2 SETTING THE SCENE 3.17.2.1 The Oceanic Lithosphere before Subduction 3.17.2.2 Continuous versus Discontinuous Reactions 3.17.2.3 Fluid Production 3.17.2.4 Fluid Availability versus Multicomponent Fluids 3.17.2.5 Real World Effects 3.17.3 DEVOLATILIZATION REGIMES IN MORB 3.17.3.1 High Dehydration Rates and Fluid Production (Typically up to 600 8C and 2.4 GPa) 3.17.3.2 Low Dehydration Rates and Little Fluid Production (2.4–10 GPa and 500– 800 8C) 3.17.3.3 Melting Regimes (650–950 8C; to 5–6 GPa) 3.17.3.3.1 Fluid-saturated (flush) melting 3.17.3.3.2 Fluid-absent melting 3.17.3.4 Dissolution Regime (.5–6 GPa) 3.17.4 HOW MUCH H2O SUBDUCTS INTO THE TRANSITION ZONE? 3.17.5 DEVOLATILIZATION IN SEDIMENTS 3.17.5.1 Pelites 3.17.5.2 Carbonates 3.17.5.3 Graywackes and Volcaniclastics 3.17.5.4 Melting of Sediments Compared to Melting of MORB 3.17.6 SERPENTINIZED PERIDOTITE 3.17.7 IMPLICATIONS FOR TRACE ELEMENTS AND AN INTEGRATED VIEW OF THE OCEANIC LITHOSPHERE 3.17.7.1 Mobile Phase Production and Trace-element Transfer 3.17.7.2 Integrating Fluid Flux over the Entire Subducted Oceanic Crust: An Example 3.17.8 CONCLUSIONS AND OUTLOOK REFERENCES

567 569 569 571 572 572 573 573 573 575 575 576 577 577 578 579 579 580 581 581 582 583 583 585 586 588

(including the supra-subduction arc magmatism), subduction itself consumes oceanic crust. This recycling process replenishes the mantle with most of the element inventory that otherwise would be, with time, strongly depleted in the mantle. Thus, subduction has a major role in

3.17.1 INTRODUCTION Subduction zones are the geotectonic settings where the Earth’s mantle is refertilized. Whereas the various magmatic geotectonic settings produce oceanic and continental crust 567

568

Generation of Mobile Components during Subduction oceanic crust to depth-equivalents of 10 GPa (Figure 1): (i) High dehydration rates at low-to-medium P, low T (,2.5 GPa, ,600 8C) where hydrous phases are abundant and dehydration reactions are often perpendicular (in P –T space) to typical subduction geotherms. All subducted lithosphere goes through this first stage. (ii) Medium to low dehydration rates at medium-to-high P, low T (2.5 –10 GPa, 500 – 850 8C) where hydrous phases are already largely reduced in volume and dehydration reactions are often subparallel to possible P – T paths. In this range fluids become increasingly rich in dissolved matter. (iii) Melting where the amount of melt depends mostly on H2O-availability and the composition of melts is, in addition, strongly pressure sensitive: (a) Flush melting (1 –4 GPa, 650– 850 8C) at temperatures between the wet granite solidus and the fluid-absent amphibole, biotite, and phengite melting curves. In this case, additional fluid is provided from underlying dehydrating lithologies.

maintaining the Earth’s magmatic environments and tectonic style over the geological history. For understanding the recycling process it is necessary to understand the reactions that occur during subduction, and within the subduction setting, in particular those that transfer material from the subducting lithosphere to the mantle wedge. Prograde metamorphism of subducting oceanic crust causes a series of mineralogical reactions that inevitably result in eclogites that may or may not contain hydrous phases and/or carbonates. An alternation of continuous and discontinuous reactions causes devolatilization, i.e., production of a fluid or, more generally, a mobile phase. The mobile phase is either a low-density fluid, a highdensity solute-rich fluid, a silicate melt, or a carbonatite melt. This contribution reviews the reaction mechanisms and conditions resulting in the generation of the various mobile phases and also examines the restite(s) that are subducted to great depths. In general, four different regimes producing a mobile phase can be recognized in the 13 12

MORB = dry hydrous phases in sediments

11

Dry

10

Phengite

Law

soni

ite

Little dehydration

6 5

ii

Chl o

Considerable dehydration

3

e

Chlorite

w La Zo

i

1

Zoisit

Amph

Amphibole

2

oid rit

4

300

400

500

600

Early slow thin

K > Na melts

iii.b

Na = K melts

iii.a

Very strong dehydration

0

e

ii

ine

git

e nt

Dissolution: solute-rich fluids

iv

en

Ser p

Ph

7

Early fast thin

Law

e Stishovit Coesite

Late fast thin

W et m Biotite elt in Amphib g ole t n e s m Fl u i d - a b elt Phe ing ngit e

Pressure (GPa)

8

son

9

te

700

800

900

Na >> K melts 1000

1100

1200

1300

Temperature (˚C)

Figure 1 Devolatilization regimes during subduction, mainly based on phase relations in MORB (compare to Figure 3). The P– T region with high dehydration rates is limited by the amphibole stability and the wet solidus, the P – T regions with less dehydration is limited by zoisite and lawsonite stabilities. H2O-saturated melting takes place in MORB and sediments at the wet solidus, fluid-absent melting to 2.5 GPa is dominated by amphibole (MORB) or biotite (pelite), and above 2.5 GPa by phengite (only in pelites). The classic melting regime is replaced by a continuous dissolution of hydrous phases in a solute-rich fluid at higher pressures (compare to Figure 4). The serpentine stability field in peridotite (dotted line) is given for reference (compare to Figure 5). The stippled arrows represent cold-to-intermediate subduction geotherms (after Kincaid and Sacks, 1997).

Setting the Scene (b) Fluid-absent melting at high temperatures ( . 800 –900 8C). At relatively low P, amphibole and biotite (1 –2.5 GPa and 1 –3 GPa, respectively) are the principal hydrous phases to melt. Adakitic magmas, i.e., slab melts (with Na . K), are likely to occur. At medium P (2.5 – 5 GPa), phengite is the principal hydrous phase to melt, and melts have K . Na. (iv) Dissolution at high P, high T ( .5 GPa, .800 8C) where the solvus between fluid and melt is closed and a solute-rich fluid might dissolve hydrous phases. In the following discussion some mechanisms for devolatilization reactions are illustrated, including real-world aspects such as chemically heterogeneous protoliths and failure to reach equilibrium (compared to a chemically homogeneous protolith and equilibrated ideal situation). We describe and quantify, as far as possible, the four different dehydration/melting regimes, principally investigating three bulk compositions, i.e., pelite, MORB, and harzburgite and H2O as the major volatile component. Subsequently, carbonates and CO2 transfer are discussed. Phase petrology is then applied to understand the behavior of trace elements. Finally, we argue for the necessity of integrating fluid/melt-producing processes over the entire oceanic lithosphere, as pressure –temperature conditions of the different lithologies within a given column of subducted lithosphere are strictly related to each other. The purpose of this chapter is to give an overview of the possible mechanisms responsible Table 1 Mineral Phengite Biotite/phlogopite Paragonite K-richterite Glaucophane – barroisite Hornblende – pargasite Lawsonite Zoisite/epidote Chloritoid Chlorite Talc Talc in Si-rich veins Serpentine Phase A Phase E ˚ phase 10 A Aragonite/calcite Dolomite Magnesite

for fluid or melt transfer into the mantle wedge, rather than to provide a catalog of minerals (see Table 1) and reactions in subducted lithosphere.

3.17.2 SETTING THE SCENE 3.17.2.1 The Oceanic Lithosphere before Subduction In the following, we summarize some aspects that are crucial for the production and quantification of mobile phases in the oceanic lithosphere (Figure 2). The unaltered igneous oceanic crust is described in detail in Chapter 3.13, the overlying sediments are described in Chapter 7.01 and alteration processes are described in Chapter 3.15. A well-stratified oceanic crust is composed of a sedimentary cover layer (including pelites, carbonates, cherts, and volcaniclastics), a basaltic layer built of pillows and sheeted dikes, and a gabbroic layer with an upper layer chemically close to MORB but a lower part mostly composed of differentiated high-magnesium gabbros, Fe – Tigabbros, troctolites, norites, and other cumulates (Nicolas, 1989). The crust is underlain by a partly serpentinized, depleted, mostly harzburgitic peridotite (Snow and Dick, 1995). All sediments have some natural porosity and begin expelling fluids during compaction. They generally remain fluid saturated during prograde metamorphism, and equilibration of major elements can be generally assumed (Carlson, 2002)

Major volatile-carrying phases in subduction zones. Chemical formula

K(Mg,Fe)0.5Al2Si3.5O10(OH)2 K(Mg,Fe)2.8Al1.4Si2.8O10(OH)2 NaAl3Si3O10(OH)2 KCa(Mg,Fe)4AlSi8O22(OH)2 NaCa(Mg,Fe)3Al3Si7O22(OH)2 – Na2(Mg,Fe)3Al2Si8O22(OH)2 amp Ca2(Mg,Fe)4Al2Si7O22(OH)2 – NaCa2(Mg,Fe)4Al3Si6O22(OH)2 law CaAl2Si2O7(OH)2·H2O zo/epi CaAl2(Al,Fe3þ)Si3O12(OH) cld (Mg,Fe)2Al4Si2O10(OH)4 chl (Fe,Mg)5Al2Si3O10(OH)8 tc (Mg,Fe)3Si4O10(OH)2 phe bt-phl par K-rich amp

serp “A” “E” 10A ara/cc dol mgs

(Mg,Fe)48Si34O85(OH)62 (Mg,Fe)7Si2O8(OH)6 (Mg,Fe)2.2Si1.1O2.8(OH)3.2 (Mg,Fe)4Si3O10(OH)2·H2O CaCO3 CaMg(CO3)2 MgCO3

þ : ,5 vol%, þþ : 5 – 20%, þþ þ: .20%.

569

Pelite Basalts Mg-gabbros Peridotite H2O (wt.%) graywackes 4.3 4.1 4.6 2.0 2.2

þþþ þþ þ 2 þ

þ þ þþ 2 þþþ

2 2 þ 2 þþþ

2 þ 2 þ 2

2.2

2

þþþ

þþþ

þ

11.2 2.0 7.5 12.5 4.8

þ þ þþ þþ þþ

þþ þþ þ þþþ þ

þþ þþ þþ þþþ þþ

2 2 2 þ þ þþ þ

12.3 11.8 11 – 18 8 – 14

2 2 2 Likely þ þ þ

2 2 2 2 þ þ þ

þ ? ? Likely þ þ þ

þþ þ þþ þ þþ þ 2 þ þ

570

Generation of Mobile Components during Subduction Active axial circulation

Passive off-axis circulation

Oxidation addition of H2O and alkali

Time Carbonate addition

CO2 0

12

Basalts Gabbros Epidosites Heat source

H2O Troctolites Harzburgites 0

4

8 12 wt.%

Figure 2 Schematic representation of hydrothermal alteration of the igneous oceanic crust. The inset to the right gives H2O and CO2 contents in a depth profile of the oceanic crust, the two horizontal bars indicate typical H2O contents in troctolites and in intensively serpentinized harzburgite.

as field evidence for prograde disequilibrium is essentially absent. Also, eclogites and blueschists having a basaltic precursor generally appear to be fully hydrated during high-pressure metamorphism, thus assuming fluid saturation and equilibrium during subduction is a reasonable simplification. In contrast, gabbros, which are only partly hydrated (in veins and adjacent alteration zones), may not experience fluid-saturated conditions and may thus fail to reach equilibrium. When pervasive alteration of coarse-grained gabbros takes place, hydration reactions are often limited to grain boundaries. Like the gabbros, serpentinization of peridotite is very heterogeneous. Veins and fractures represent important infiltration pathways for fluids and zones of intensive serpentinization probably alternate with very weakly altered peridotite. Mammerickx (1989) estimated that 20% of the Pacific Ocean floor is affected by fracturing. For both peridotite and gabbros it remains difficult to provide quantitative estimates of the volume affected by hydration. In addition, the amount of hydration/carbonatization also depends on the spreading velocity at the midocean ridges (MORs); oceanic lithosphere produced at slow-spreading ridges appears to undergo higher amounts of hydration than that produced at fast-spreading ridges. Alteration of the igneous parts of the oceanic crust also forms veins containing carbonates, and although pervasive carbonatization is rare, carbonate contents might be locally high in veins and their immediate surroundings. The well-stratified oceanic-crust paradigm applies to the circum-Pacific, where 53% of the

total length of present Earth oceanic-crust subduction takes place (own compilation). Fast spreading Pacific-type oceanic crust, however, is not found in the Atlantic and Indian oceans where 4% and 16%, respectively, of present subduction occurs (the remainder being subduction of oceanic crust from comparatively small basins). The slowspreading Atlantic has an oceanic lithosphere that is more complex structurally, and in which lateral heterogeneity is an important feature (Gente et al., 1995; Cannat, 1993). Large transforms offer a preferential site for the emplacement of ultramafics in the shallowest portions of the lithosphere (Constantin, 1999). In the Atlantic it is common to have serpentinites cropping out on the ocean floor; such bodies are rare in the Pacific. The Indian Ocean has a crustal architecture intermediate between the Atlantic and Pacific oceans, it currently subducts in the Macran arc and from Burma to Sunda. Thus, caution is necessary when applying a simple layered oceanic-crust model to infer the geometry of subducted lithologies in the past and in some present subduction zones as, for example, S-Sandwich, Antilles, or Sumatra. Determining the state of the oceanic lithosphere before subduction is crucial for understanding how the different lithologies evolve. Most of the alteration takes place close to the mid-ocean ridges, however, recent reports show that some hydrothermal activity takes place also in older oceanic crust away from the ridge axis (Kelley et al., 2001). Fracturing at the seafloor and fractures caused by bending of the subducting lithosphere at the trench provide pathways for fluids and hence, further possibilities of hydration. Once the oceanic

Setting the Scene

both the sedimentary and mafic layers of the oceanic crust. Discontinuous reactions signify the appearance and/or disappearance of phases, whereas continuous reactions only change composition(s) and proportions of the phases already present. Extensive solid solutions (amphiboles, micas, pyroxenes, garnet) result in continuous reactions that release fluids over several tens of kilometers depths. This can be exemplified by the disappearance reactions of amphibole (Figure 3(a)) within the amphibole-eclogite facies (Poli, 1993). At the minimum pressure necessary for the formation of omphacite in basaltic bulk compositions (1.5 GPa at 650 8C), more than 50% amphibole remains in the eclogitic assemblage omphacite – garnet – amphibole – epidote – quartz ^ paragonite. Within this assemblage, amphibole decomposes progressively, mainly forming omphacite and garnet, until 22% amphibole remains at its upper pressure stability at ,2.3 GPa. Within this pressure interval, amphibole composition changes from calcic and tschermakite-rich to sodic –calcic and barroisite-rich, and the continuous reaction from 1.5 GPa to 2.2 GPa produces more fluid (i.e., 0.7 wt.% H2O) than the discontinuous terminal amphibolebreakdown at 2.2 GPa (i.e., 0.4 wt.% H 2O), which results in chloritoid as an additional hydrous phase. As in mafic compositions, amphibole decomposition and abundance in peridotites are controlled by continuous reactions that decompose 50 –70% of the amphibole present at low pressures before its terminal-pressure breakdown at 2.5– 3.0 GPa (e.g., Niida and Green, 1999). These experimental studies demonstrate that continuous reactions and discontinuous terminal breakdown

lithosphere is subducted, continued alteration of the upper layers results from any fluid or melt expelled from deeper levels. Thus, during ongoing subduction the igneous oceanic crust will interact with fluids passing through it from the serpentinized peridotite and the sediments will interact with fluids coming from the basaltic and the serpentinized peridotite layers (Gieskes et al., 1990). The building blocks of the oceanic crust define the major lithologies involved in devolatilization and melting. To a first approximation, the major players are pelites, basalts, and serpentinized harzburgites. We consider here pelites and carbonates as the main components of the sedimentary layer. From a phase petrological point of view, graywackes and volcaniclastic sediments of broadly andesitic to dacitic composition can be regarded as intermediate between pelites and basalts. In terms of major elements and dehydration, the sediment layer is thin and volumetrically unimportant (with the exception of K2O). Its major geochemical significance is as a source of incompatible trace elements, which are concentrated in the sediments, and also in contributing to carbonate cycling.

3.17.2.2 Continuous versus Discontinuous Reactions The production of fluids and melts during subduction (the latter occurring in strictly prograde, relatively low-temperature –high-pressure conditions compared to typical crustal anatepis) is dominated by a succession of continuous and discontinuous reactions that are of comparatively equal importance in terms of fluid productivity in MORB+H2O gar

cld

650 ˚C epi

cpx

gar

571

MORB+H2O+CO2

2.4 2.2

665–680 ˚C

mag

cpx

0

40

plg 60

Abundance (vol.%)

epi 80

1.4 100

par

dol

amp

gar

dol

amp 20

qtz

1.6

amp gar

P (GPa)

par

epi

qtz

2.0

amp

0

epi 20

plg

amp 40

60

80

100

Abundance (vol.%)

Figure 3 Continuous versus discontinuous reactions contributing to the disappearance of amphibole in (a) H2Osaturated, CO2-free MORB (Poli, 1993) and (b) MORB saturated in an H2O – CO2-fluid (Molina and Poli, 2000). Modal proportions of minerals in vol.% as a function of pressure, H2O-contents for each mineral given in Table 1. Note that in both cases continuous reactions produce at least as much fluid (i.e., decompose as much amphibole) as discontinuous reactions (gray bars). In the CO2-bearing system, amphibole abundances are similar to the CO2-free system, epidote is replaced by dolomite at 1.5 GPa, and paragonite is more abundant as the appearance of omphacite is retarded to .2.1 GPa. See Table 1 for mineral abbreviations.

572

Generation of Mobile Components during Subduction

reactions are equally important in fluid generation in subduction zones, in contrast to earlier models (e.g., Tatsumi, 1986), which related fluid flow and the position of the volcanic front within a subduction zone directly to amphibole breaking down exclusively at a given depth. In sediments and the mafic portion of the subducted crust, all of the reactions involving hydrous phases and carbonates involve solid solutions whose compositions depend on the bulk composition, in addition to pressure and temperature. Different bulk compositions cause different phase compositions and thus cause reactions to shift in P– T space, i.e., to start shallower or deeper. In peridotites, amphibole and to some extent chlorite are controlled by continuous and discontinuous reactions; however, the other volumetrically important hydrous phases (e.g., brucite, serpentine, talc, and “phase A”) in altered harzburgites display a relatively restricted compositional range, at least compared to those present in mafic eclogites. As a result, breakdown reactions of hydrous phases in harzburgites are dominated (in a first approximation) by discontinuous reactions, and take place over a restricted depth range of only a few kilometers.

3.17.2.3 Fluid Production There has been a misconception in the literature that hydrous phases break down in the absence of a free fluid phase and thus the stability fields of hydrous phases are not very relevant, as they decompose when the fluid leaves the rock. This is fundamentally wrong: H2O and CO2 are chemical species just like any other species (e.g., SiO2, Al2O3, MgO). The only difference is that the phase corresponding to the composition of such chemical species (i.e., H2O, CO2, etc.) happens to have a physical state (i.e., fluid) that is different from other phases on composition of chemical species (e.g., quartz, corundum, periclase). It should be remembered that the thermodynamical treatment of all these phases is identical for the entropy and enthalpy terms and that they only differ for the P –V – T (pressure – volume –temperature) relation adopted. Just as SiO2 saturation is not a prerequisite for the stability of olivine or enstatite, so H2O saturation is not required for the stability of hydrous phases. In fact, for a given bulk composition it is quite possible (when following a suitable P – T trajectory) to pass from a fluid-absent to a fluid-present regime and back again to a fluid-absent regime. A fluid is only produced if a given rock volume is already completely hydrated (fluid saturated). If fluid saturation is not realized at the beginning of subduction, a number of fluid-absent reactions will take place. These reactions are of the type

A þ V1 ¼ B þ V2 (where A, B are volatile free phases and V 1, V 2 are hydrous phases or carbonates), involve hydrates and/or carbonates and change the mineralogy of a rock volume according to the stability fields of the minerals, but do not liberate a fluid. Prograde subduction zone metamorphism (as is true for any type of prograde metamorphism) generally reduces the amount of H2O that can be stored in hydrous minerals with depth. Thus, almost any part of the oceanic crust sooner or later becomes fluid saturated. In an equilibrium situation, the volatile content bound in hydrous phases and carbonates remains constant until fluid saturation occurs. Either continuous or discontinuous reactions may lead to fluid saturation in a rock. The point at which this occurs depends on initial water content, and pressure and temperature, and somewhat counter-intuitively, initial low water contents do not cause early complete dehydration, but delay the onset of fluid production to high pressures. Due to heterogeneous alteration (and thus varying initial H2O and CO2 contents) there is a wide depth range over which different volumes of the oceanic lithosphere become fluid saturated. A second complication arises from the scale at which equilibrium is effective. A few grains may locally form a fluid-saturated environment, but it is questionable whether the fluid produced on a local grain scale is able to escape. Field evidence argues for equilibration of fluids in eclogites on a centimeter to meter scale until they are able to collect in veins where these fluids might escape (Philippot, 1993; Widmer and Thompson, 2001; Zack et al., 2001; see Chapter 3.06), possibly through fractures in the overlying lithologies without affecting them much, possibly through pervasive infiltration. Thus, considerable uncertainties regarding the quantification of fluidor melt-producing processes result from the scale of equilibration and the way fluid migrates (Austrheim and Engvik, 1997) through the overlying layers. The ways fluids migrate also influence the phase equilibria and reactions taking place. Most reactions during subduction occur in response to an increase in pressure and temperature, where fluid is produced and expelled. However, the ascending fluid may change the fluid composition in the overlying layers and thus cause reactions. 3.17.2.4 Fluid Availability versus Multicomponent Fluids In calculations of phase equilibria that appear to be H2O undersaturated, it is commonly assumed that a fluid phase is present and H2O activity is lowered by CO2 in the fluid. The latter is not necessarily true, and this section examines the differences between the limited availability

Devolatilization Regimes in MORB of an aqueous fluid compared to the unlimited availability of a mixed H2O – CO2 fluid. Note that both cases are described by the term “H2O undersaturated.” This can be illustrated by a natural example. In the coarse-grained Allanin magnesium-gabbro, infiltration of fluid caused the formation of reaction rims around olivine (Chinner and Dixon, 1973). The succession is olivine ! anthophyllite (2 wt.% H2O) ! talc (4wt.% H2O þ kyanite ! chloritoid (8wt.% H2Oþ talc þ kyanite. This reaction rim is H2O undersaturated, and the succession of mineral assemblages corresponds to an increase of H2O content towards the rim and can only be modeled by an increase in the availability of water towards the rim. The H2O-undersaturated character of the inner rim zones does not necessitate (or justify) a CO2 component in the fluid, but rather reflects limited availability of an H2O fluid. Thermodynamic calculations based on measured compositions of solid phases (as commonly performed) result in an evaluation of the chemical potential of H2O in these phases (mH2O). Under equilibrium, the chemical potential of H2O is equal in all phases and only if some additional constraint implies the presence of a fluid phase (e.g., fluid inclusions), the composition of this fluid phase can be calculated from P, T, and mH2O. This can be illustrated in the simple system CaO – Al2O3 – SiO2 – H2O – CO2 where a given chemical potential of H2O at a given P and T (4 GPa, 600 8C) may correspond to a fluid-absent situation (with lawsonite þ zoisite þ aragonite þ coesite þ kyanite present) or to a situation with a mixed fluid phase (with aragonite þ coesite þ kyanite þ fluid present) (Poli and Schmidt, 1998). The actual consequence of H2O undersaturation in the context discussed here is that mixed fluid phases, i.e., solutions, will shift phase equilibria in P – T –X space, whereas the presence or absence of a single component fluid (e.g., a purely aqueous fluid) determines whether or not a reaction takes place, without changing the reaction’s position in P – T space.

3.17.2.5 Real World Effects The compilation of available phase equilibria aims at understanding equilibrium situations in typical (and homogeneous) average bulk compositions. However, different real-world effects with amplitudes that may depend on the rock type, are to be expected. In the real world, the following factors contribute to the continuous character of the devolatilization signal from the downgoing slab: highly variable bulk compositions in the sedimentary and gabbroic layers as well as possibly different degrees of depletion (caused by different amounts

573

of melt extracted) in the hydrated peridotitic layer (Constantin, 1999); heterogeneous distribution of carbonates versus hydrous minerals resulting in an inhomogeneous XCO2 in the fluid phase (Gillis and Robinson, 1990); large temperature gradients within the subducting lithosphere (Kincaid and Sacks, 1997), and finally, possible kinetic effects that inhibit reactions and thus widen reaction zones, which are related both to the effects of fluid availability and deformation history (Austrheim and Engvik, 1997; Molina et al., 2002) and to thermal retardation of sluggish solid – solid transformations (Schmeling et al., 1999). All of the above effects lead to smearing out the release of a fluid/melt pulse over a broader depth range. In contrast, the single focusing mechanism for fluid/melt flow is mechanical: fluids/melts may ascend through channelized flow, hydrofractures or pre-existing fractures, which may focus a broadly distributed fluid into distinct pathways. The relative importance of these mechanisms is largely conjectural, and whether macroscopical fluid/melt focusing is achieved depends on their extent and spacing.

3.17.3 DEVOLATILIZATION REGIMES IN MORB Based on phase relations in H2O-bearing MORBs, four distinct P– T regions with characteristic mobile-phase production mechanisms can be identified and are discussed sequentially.

3.17.3.1 High Dehydration Rates and Fluid Production (Typically up to 600 8C and 2.4 GPa) Once the oceanic crust starts subducting, most of its remnant porosity will be immediately lost by compaction and its pore fluids get expelled. At this stage, zeolites, pumpelleyite, and prehnite are the major H2O-bearing minerals and H2O contents stored in hydrous minerals amount to 8 –9 wt.% H2O in the bulk rock (Peacock, 1993). Beyond depths of ,15 km the oceanic crust enters into the blueschist facies in which the major hydrous minerals are chlorite, sodium-rich, calcium-poor amphiboles (glaucophane to barroisite), phengite (white mica), lawsonite or zoisite, and paragonite (e.g., Sorensen, 1986). Water contents of the bulk rock at the beginning of the blueschist facies are ,6 wt.% (Figure 4). Initially abundant chlorite has high H2O contents (12 wt.% H2O) and decomposes completely in the depth range to 70 km through various continuous and discontinuous reactions. Lawsonite (11 wt.% H 2O) has a maximum abundance of 25 vol.% at the onset of blueschistfacies metamorphism and decreases to ,10 vol.%

574

Generation of Mobile Components during Subduction 10 Law

9

MORB + H 2O

soni

K-holla ndite Phengite

te 0.1

0.1

8 e Stishovit coesite

7

0.3

0.1

2 1

0 300

0.5

du So li

e git

0.45 0.9 1.1 0.5 0.9 0.1 w 0.6 zo 1.0 La o 0.5 Z Chloritoid 0.8 cld 0.45 para amph Amphibole 1.1 2.3 te ni 1.5 ara so ite p w s 1.8 3.8 La Zoi 2.3-1.5 i 3.1 cpx 1.4 ep 5.4-4.1 1.6-1.8 6.0 1.3 Garnet 3.0 1.2 5.4 1.4 1.5 3.5 0.9 2.7 4.4 ote Solidu Epid s 1.2

Chlorite

3

id ito

4

Law

5

Ph en

son ite

s

0.4

lor Ch

Pressure (GPa)

6

400

500

600 700 Temperature (ºC)

800

900

Figure 4 Major phase stability boundaries in H2O-saturated MORB and H2O contents (numbers in wt.%) stored in hydrous phases (Schmidt and Poli, 1998). Data below 550 8C are based on natural blueschists and greenschists, all other data are based on experiments (compare with Figure 1).

at the amphibole-out reaction. Anhydrous minerals typically comprise 5– 25 vol.% (e.g., Thurston, 1985; Okay, 1980) at 5– 10 km depth and some 100 8C, grow in volume to ,50 vol.% just before the amphibole-out reaction, and increase to .70 vol.% at pressures beyond amphibole stability. As a consequence, dehydration rates are considerable within this P– T regime, where a fully hydrated MORB loses ,4 – 6 wt.% H2O when passing through the blueschist stage within the fore-arc region. At pressures of 2.2– 2.4 GPa, i.e., the maximum pressure stability of amphibole in MORB (65 – 70 km), dehydration reactions are numerous and their orientations in P – T-space are mostly oblique to a typical subduction-type P – T path, resulting in high dehydration rates. Although fully hydrated oceanic crust loses about two-thirds of its initial water content in this interval, this fluid will either serpentinize the cold corner of the mantle wedge or eventually pass, through veins, to the ocean floor in the fore-arc region (e.g., Mariana arc, Fryer et al., 1999). The high dehydration rates in conjunction with a—close to the trench—thin mantle wedge should lead to rapid, full serpentinization of the cold corner (i.e., within 0.8– 3.0 Ma) if fluids pass pervasively through the

mantle wedge (see also Gerya et al., 2002). As a serpentinized cold corner has a much lower viscosity than anhydrous mantle, the serpentinized cold corner may develop a counter-convection cell and stagnate in its corner, whereas the rest of the mantle wedge undergoes large-scale convection (examples of cold corner counter-flow are found in Davies and Stevenson, 1992). Most information concerning blueschists is obtained from studying natural occurrences; it is only at and beyond the chlorite- and amphibole-out reactions that reaction rates are sufficiently high to study these systems experimentally. Thermodynamic modeling of phase relations in MORB is generally accurate, however, it becomes difficult when incorporating the amphibole solid solution. Both calculated amounts and stabilities of amphiboles are far from what is observed in nature and experiments (e.g., the pressure stability of amphiboles in MORB is calculated to 4– 5 GPa instead of the 2.4 –2.8 GPa observed in experiments and deduced from natural eclogites), leaving us to rely almost entirely on natural occurrences to deduce relevant dehydration reactions. The study of blueschists also demonstrates that many of their minerals contain significant

Devolatilization Regimes in MORB 3þ

amounts of Fe (e.g., Maruyama et al., 1986; Brown and Forbes, 1986). The inhomogeneous degrees of oxidation within the altered oceanic crust add another compositional variable, shift reactions in P – T-space, and complicate the geochemistry of trace elements with variable oxidation state (e.g., uranium). The fraction of Fe3þ/Fetot is generally highest at the onset of subduction-zone metamorphism and decreases with increasing grade. 3.17.3.2 Low Dehydration Rates and Little Fluid Production (2.4–10 GPa and 500–800 8C) Until the late 1980s it was believed that oceanic crust is fully dehydrated after the pressure-induced amphibole breakdown. However, experiments (Poli, 1993; Pawley and Holloway, 1993; Poli and Schmidt, 1995; Schmidt and Poli, 1998; Okamoto and Maruyama, 1999; Forneris and Holloway, 2001) and natural occurrences of epidote/zoisite, lawsonite, talc, chloritoid, phengite, staurolite, OH-rich topaz, and many other minor hydrous minerals in eclogites that have already lost their amphibole, clearly demonstrate that dehydration continues above 2.4 GPa. In particular, since the early 1990s, the frenetic chase after coesite- and diamond-bearing hydrated eclogites showed that most of the above hydrous minerals are stable at pressures beyond 2.5 GPa (e.g., the coesite-lawsonite bearing eclogites found in xenoliths of Colorado Plateau, Helmstaedt and Schulze, 1988; Usui et al., 2003). Generally, the natural occurrences of coesite- and sometimes diamond-bearing hydrous eclogites (the most famous being Dora Maira, Norway, Dabie Shan, and the Kokchetav massif, see Liou et al. (1998) and Chapter 3.09 for a concise review) are in good agreement with experimental findings, and the discussion below is based on experimental P –T conditions and mineral abundances. Above 2.4 GPa, a maximum of 1.5 wt.% H2O remains stored in hydrous phases of the oceanic crust. The major water hosts are lawsonite, zoisite, chloritoid, talc, and phengite. Lawsonite is the most water-rich of these phases (12 wt.% H2O) and may host .50% of the water present; however it is restricted to relatively low temperatures (Figure 4). Zoisite is stable to ,3.3 GPa and occurs at the wet solidus until 3 GPa. Talc is only a minor phase in MORB but becomes abundant in bulk compositions with high XMg, i.e., magnesium gabbros. Phengite has an important role, as virtually any K2O present in the oceanic crust at pressures below 5 GPa will be stored in phengite, irrespective of the bulk composition (Schmidt, 1996). Above this pressure, potassium also enters into clinopyroxene (Okamoto and Maruyama, 1998; Schmidt and Poli, 1998) and the relative

575

amounts of phengite, clinopyroxene, and the coexisting fluid will depend upon the quantity of dissolved potassium and its dissolution rate into the fluid. Above 2.4 GPa continuous reactions dominate over a few discontinuous reactions and most (including the wet solidus) are subparallel to typical subduction P –T paths. As a consequence, a particular rock volume crosses reactions over a wide depth interval. Thus, dehydration rates are rather low, and the (post-amphibole) remnant 1.5 wt.% of H2O is lost over a wide depth range. The pressure – temperature range that defines this “low dehydration rate” regime corresponds to the depth interval of the subarc region. The hydrous phases in oceanic crust with the widest pressure stability range are lawsonite and phengite, which both reach into the stishovite field. The breakdown of these two phases corresponds to the end of the major devolatilization history of oceanic crust. Because many reactions are parallel to typical P – T paths during subduction, effective H2O contents, mineral assemblages, and mineral compositions are fairly sensitive to temperature. A K2O-free MORB passing 600 – 650 8C at a depth of 100 km has 1 wt.% H2O stored in lawsonite and chloritoid and will arrive at 200 km with ,0.4 wt.% H2O in lawsonite. Passing 100 km at a temperature of 700 8C causes the loss of the last hydrous phase (zoisite) near this depth and our MORB volume becomes a dry eclogite and thus irrelevant for any further production of mobile phases. Now, if the 100 km depth is achieved at a temperature of 750 8C, a MORB is already above the wet solidus and any fluid infiltrating from below should cause partial melting. As is evident from these considerations, the major difficulty in predicting what type and quantity of mobile phase is produced stems from the uncertainty in temperature distribution in the subducting oceanic crust. Unfortunately, a difference in temperature of 100 8C has a significant impact. 3.17.3.3 Melting Regimes (650–950 8C; to 5–6 GPa) Melting of the oceanic crust is evidenced in so-called adakites which form volcanic suites of andesitic to rhyodacitic composition (Kay, 1978; Drummond et al., 1996). These are interpreted to be partial melts of the oceanic crust more or less modified (mainly magnesium enriched) during their ascent through the mantle wedge. Adakitic suites are distinguished from melts originating from a fluid-metasomatized mantle wedge by distinctive trace-element ratios (i.e., high La/Yb and Sr/Y, review by Martin, 1999), by anomalously high XMg values and MgO concentrations at comparative SiO2 contents (thought to be acquired when slab melts absorb olivine on

Generation of Mobile Components during Subduction

3.17.3.3.1

Fluid-saturated ( flush) melting

Fluid-saturated melting of basaltic crust begins at temperatures of ,650 8C at 1.5 GPa to ,750 8C at 3 GPa (Figure 5). It should be noted that

the wet-solidus temperature is elevated by at least 100– 200 8C, if the fluid is in equilibrium with carbonates (an XCO2 of 0.3 – 0.6 would be expected in the appropriate P – T range). Although a small quantity of free fluid (,0.1 vol.%) is likely to be present in any lithology affected by dehydration, this would not be sufficient to produce a significant melt portion through fluidsaturated melting. However, fluid-saturated melting at relatively low temperatures could be MORB pelite

K-holla nd Phengit ite e

10

9

Law

K-hollan d Phengiteite

son

ite

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8

7

5

en

4

Z oi

700

800

ut

Bio Me tite lt

t Mel

1 600

e Phengit Bt-out Am -o ite ng elt e Ph M Garnet

site-out Melt

Zoisite

2 Cpx Ab

Coesite Quartz

Ph

3

ION

Phe

6

LUT

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Ph Lla wso en git nite e + gi cp tex ou Me jad + t H lt 2O

O DISS

Melt phibole-out

their way through the mantle wedge) and by the absence of volcanic rocks more basic than andesite. Melting of subducted crust was probably predominant on the Archean Earth (Martin, 1999) where mantle temperatures were higher and oceanic plates probably were thinner and hotter. There is consensus that on the modern Earth dehydration dominates in subducting lithosphere. However, when young, hot crust is subducted, or when the thermal field within a subduction zone is disturbed at plate boundaries (i.e., at the lateral end of a subducting plate) or due to flat slabs caused by ridge subduction, melting of the oceanic crust may occur. Melting of subducted crust can occur under fluid-saturated or fluid-absent conditions. The principal hydrous phases involved are phengite, biotite, epidote/zoisite, and amphibole. If melting of the oceanic crust occurs at pressures below 2.5 GPa, the oceanic crust will go through a greenschist and epidote-amphibolite facies stage at low pressures (both stages not discussed above). The compositions of slab melts principally depend on the bulk composition, the degree of melting and the pressure of melting. However, it is generally true that at low- to moderate-melt ratios, where cpx is a major residual mineral, Na/K ratios mostly depend on the pressure of melting. At any pressure, potassium strongly partitions into the melt. However, Nacpx/melt partition coefficients are small (,0.5) below 2 GPa, close to unity around 3 GPa and increase to 3 – 5 at 4 GPa. The latter corresponds to ,50– 70% jadeite component in residual clinopyroxene, which retains most of the Na2O, but leaves all of the K 2O in the melt. As a consequence, melts from amphibole-dominated fluid-absent melting have Na . K, whereas melts resulting from fluid-present or fluid-absent phengite melting at higher pressures are strongly potassic peraluminous granites (with K . Na). The latter, which would migrate into the mantle wedge at .70 km depth are not observed at the surface. This is either because they never form or because it is virtually impossible for them to traverse the thick mantle wedge and retain their peculiar chemistry, which is far out of equilibrium with peridotite. Absorption reactions with mantle minerals would largely modify the majorelement composition of these melts, although a sediment-melt contribution is sometimes postulated in arc magmas on the basis of certain trace elements (see Section 3.17.5).

Pressure (GPa)

576

900

1000

1100

Temperature (˚C)

Figure 5 Compilation of melting reactions in MORB and average pelite. Black lines represent reactions with similar P – T locations in both bulk compositions, blue lines: MORB, red lines: pelite. The red and blue areas are P – T fields of biotite-dominated and amphiboledominated fluid-absent melting respectively. At subduction-zone melting pressures (i.e., .1 GPa), the first fluid-absent melt in pelite appears at the phengiteout reaction and in MORB at the zoisite/epidote-out reaction. The bold lines represent the H2O-saturated solidus, which involves amphibole or biotite in addition to phengite þ cpx þ coesite/quartz below 2.5 GPa and plagioclase instead of cpx below 1.5– 1.7 GPa. Upto 5 – 6 GPa classical melting occurs where a solidus separates a crystal þ fluid field from a melt-present field (the fluid having distinctively higher H2O contents than the melt). A continuous increase of solute in the fluid is observed at high pressures (.5.5 GPa) and at least the K-bearing phases (phengite or at higher pressures K-hollandite) dissolve in a fluid of continuously evolving composition.

Devolatilization Regimes in MORB

577

achieved in subducted crust through flush melting, i.e., by addition of aqueous fluid from below. In this case, the melt productivity depends on the amount of fluid added to the system. The possibilities to obtain volumetrically significant melt fractions in (i) MORB through flushing with fluid originating from dehydration of serpentinized peridotite situated below the MORB or (ii) sediments through flushing with fluids from underlying MORB or serpentinized peridotites are discussed in Section 3.17.5.4.

layer (800 –900 8C) are such that the hydrated peridotite below would reach the serpentine stability limit, thus possibly providing a significant amount of fluid. This fluid would strongly increase the melt fractions in amphibolites and metapelites and might be required to produce a sufficient quantity of primary melts that then give rise to adakite suites.

3.17.3.3.2

The typical concept of low-density H2O –CO2 fluids (with small to moderate amounts of solute) in the sub-solidus and high-density silicate liquids (with typically 1– 15 wt.% H2O dissolved) above the solidus does not apply to the subduction environment at pressures above 5 – 6 GPa. At higher pressures, a chemical continuum between fluids and melts exists (Boettcher and Wyllie, 1969) and, depending on fluid – rock ratios, a continuous dissolution process leaches hydrophile species out of sediments, basalts, and serpentine. Considerable attention has focused on the role of high-pressure fluids/melts that exist beyond the second critical end point (Ricci, 1951). At crustal pressures and temperatures, a large miscibility gap exists between a low-density aqueous (already supercritical) fluid and a high-density hydrous silicate melt. Consequently, a fluid and a melt have quite distinct compositions and physical properties (e.g., viscosity, compressibility). This is reflected in the common terminology of dehydration/hydration versus melting reactions. However, at higher pressures, the solubilities of silicates or silicate components in fluid and of H2O in silicate melts increase. Beyond a certain pressure (depending on the chemical system, between 1.5 GPa and .12 GPa), the miscibility gap between classical fluid and melt shrinks, intersects the solidus (this exact locus is defined as the second critical end point) and a chemical continuum between the two extremes, a dry silicate melt and a pure H2O-fluid, is possible (see Stalder et al., 2000, figure 1). This implies that a continuum exists between the physical properties of former melt and fluid; it does not imply that a mobile phase with melt chemistry and physical properties characteristic of silicate melts may not exist (or with fluid chemistry and properties). The term supercritical fluid is not distinct or adequate enough to describe this phase, as any aqueous/carbonic fluid above the first critical end point (at a few hundred bars and degree centigrade) is already supercritical. At the conditions described above, the wet solidus vanishes, the concept of melting loses its definition, and solid assemblages continuously dissolve in first, a volatile-rich, and with increasing temperature, silicate-rich nonsolid phase.

Fluid-absent melting

Most adakite suites contain andesites that are consistent with fluid-absent melting of a basaltic source (but not with melting of mica-dominated sediments only). Fluid-absent melting is defined as the production of a silicate melt with lowwater contents from an assemblage that contains hydrous phases but no free fluid phase. At pressures below 2.5 GPa, both metapelites and metabasalts contain a pair of hydrous phases with equal water contents (Figure 5): in pelites, phengite (4.3 wt.% H2O) has a lower modal abundance than biotite (4.1 wt.% H2O) and also a 100 – 150 8C lower melting temperature than biotite. In basaltic compositions, epidote/zoisite (2.0 wt.% H 2O, typically 10 vol.%) melts ,100 8C below amphibole (2.2 wt.% H2O, typically 50 – 30 vol.%, decreasing with increasing pressure) (Vielzeuf and Schmidt, 2001, and references therein). In both cases, the first fluidabsent melts are produced through the volumetrically less-important phase, i.e., phengite in metapelites and epidote/zoisite in metabasalts. These first melts are thought to be dacitic and probably amount to less than 10 vol.%. In the absence of detailed experimental investigations (at fluid-absent conditions, experimental reaction rates are too sluggish below ,800 8C and equilibration does not take place for the fluidabsent melting of epidote and phengite below 2.5 – 3 GPa), it is likely that the primary melts of adakites result from 20 – 35% amphiboledominated fluid-absent melting of MORB, geochemical modeling is consistent with a garnet þ cpx ^ amph residue (Martin, 1999). Nevertheless, the temperature distribution in oceanic crust is such that subducted sediments are at higher temperatures than the igneous oceanic crust. Thus, if amphibolites are melted, the overlying sediments must also melt through mica-dominated fluid-absent melting. This complexity needs to be taken into account in geochemical slab-melting models. Fluid-absent melting of amphibolite or metapelite cannot be fully described in a closed system or a system open only to melt extraction. The temperatures necessary for fluid-absent melting of the MORB

3.17.3.4 Dissolution Regime (>5–6 GPa)

578

Generation of Mobile Components during Subduction

The exact pressures (at temperatures relevant for subduction) for the disappearance of the wet solidus are somewhat uncertain and strongly depend on the chemical system. In the simple system SiO2 –H2O, the second critical end point is only at 1 GPa, 1,100 8C (Kennedy et al., 1962), in albite-H2O it moves to ,1.5 GPa (Stalder et al., 2000), in CaO – SiO2 – H2O – CO2 it occurs at 3.2 GPa, 500 8C (Boettcher and Wyllie, 1969), whereas in the model ultramafic system MgO –SiO2 – H2O (MSH) the solidus terminates at ,12 GPa, 1,100 8C (Stalder et al., 2001). In potassium-enriched MORB, and in graywackes and pelites, the miscibility gap closes somewhere between 4 GPa and 6.5 GPa (Schmidt and Vielzeuf, 2001). At 4 GPa, a classical sequence of melting reactions and quenched melts are observed in experiments, whereas at 6.5 GPa, initially abundant phengite becomes less and less abundant until it disappears and a mostly K2O – Al2O3 – SiO2-bearing fluid-quench precipitate becomes more and more abundant at grain boundaries. This finding is in agreement with an experimental study on carbonaceous pelites (Domanik and Holloway, 2000), where phengite disappears in a similar fashion at 6.5 GPa and 8 GPa. Thus, there is little doubt that at high pressures the solvus closes and a chemical continuum is realized. Estimates of the amount of solute in such fluids above 6– 7 GPa range from 50 wt.% to 70 wt.% (a molar H2O : K2O ratio slightly above unity is in fact sufficient for completely dissolving micas), being close to the solubility of H2O in watersaturated melts near 3 GPa (solubilities of H2O in natural silicate melts at higher pressures are unknown). Due to the high solubility of K2O (and probably other components that are less soluble at low pressures), it is doubtful that much of the subducted potassium and related trace elements could reach great depths. In fact, most of the potassium is likely to be transferred to the overlying mantle wedge, where in a MgO-rich, SiO2-poor chemical environment it may reprecipitate in phlogopite or potassium-richterite (potassic amphibole) (Sudo and Tatsumi, 1990; Konzett and Ulmer, 1999; Trønnes, 2002) and then be dragged to greater depths in the mantle directly overlying the subducted slab. The little potassium remaining in the oceanic crust after dehydration and leaching is stored in the anhydrous phase potassium-hollandite (KAlSi3O8), which is stable to .25 GPa.

3.17.4 HOW MUCH H2O SUBDUCTS INTO THE TRANSITION ZONE? It is well known that refertilization of the mantle takes place through transfer of fluids or melts from

the subducting lithosphere to the overlying mantle wedge, as described above. A portion of the elements transferred into the mantle wedge is partitioned into partial melts, which ultimately form arc volcanism and thus do not refertilize the deep mantle. The residue of oceanic crust remaining from this process is subducted to depths .300 km where it may ultimately be mechanically mixed with mantle material (Allegre and Turcotte, 1986). Dixon et al. (2002) deduced from the geochemistry of ocean-island basalts, which are widely believed to contain some recycled oceanic crust in their sources, that dehydration is .92% efficient during subduction. In order to evaluate how much water is subducted to great depths in the mantle, it is necessary to (i) determine the amount of H2O stored in peridotite after pressure induced decomposition of serpentine and (ii) understand the state of the oceanic crust at pressures just above the phengite breakdown reaction. Any H2O stored in oceanic peridotites that pass beyond 220 km depth and any oceanic crust that passes beyond 300 km depth is unlikely to be mobilized within the direct subduction context and will participate in mechanical mixing with deep mantle. Only a few hydrous phases may exist beyond 220 km and 300 km depth in peridotite and oceanic crust, respectively, and temperature stabilities for these phases (“phase A,” “phase E,” and “phase D” in peridotite: Angel et al. (2001), Ulmer and Trommsdorff (1999), Frost (1999), and Ohtani et al. (2000); and the hydrous aluminosilicates topaz-OH and “phase egg” in aluminous sediments, Ono (1998)) increase faster than any subduction geotherm, at least to 20 GPa. Furthermore, hydrous phases become much less important at pressures beyond 10 GPa: the volumetrically dominant nominally anhydrous phases (NAMs) such as olivine, wadsleyite, garnet, and cpx (3,000 ppm OH in natural cpx from 6 GPa, 1,000 8C; Katayama and Nakashima (2003); up to 3.3 wt.% H2O in wadsleyite, Kohlstedt et al. (1996)) dissolve considerable amounts of hydrogen at these pressures and become the principal hydrogen reservoirs at greater depths. Whether a significant amount of water is subducted beyond 200 km in peridotitic compositions depends on the exact P–T path. As can be seen in Figure 6, any serpentine-bearing peridotite descending along geotherms that are cooler than 580 8C at 6 GPa (termed the choke point by Kawamoto et al. (1995)) will conserve H2O and form phase A (and subsequently phase E and phase D). In oceanic lithosphere subducting along geotherms that pass between 580 8C, 6 GPa and ˚ phase forms upon serpentine 720 8C, 7 GPa the 10 A breakdown (Fumagalli and Poli, 1999; Fumagalli et al., 2001), and holds 0.6 wt.% H2O in the peridotite. Subsequent entering into the stability field of phase A at greater depths will not lead to

Devolatilization in Sediments

579

8

240 ” “A H

” “A 2O AH 10

200

10 A

4.6 5.1

10A H2O

chl

dus l ch

O H2

1.0 1.8 chl

serp chl

100 H2 O

chl amp

mp

pa

ser

4.8

se p H 2O ser O amp tc H 2

500

5.0 1.6 2.3

600

tc

a

tc

amp sp gar

amp

Melt 0.2 0.4

50

amp

l

hl rp c

mp

0.2 0.4

amp chl 0.8 1.9 ch

1

150 Wet s oli

se

5.9 7.8

2

0 400

cpx opx ol gar

0.5 0.8

Depth (km)

“A” serp H 2O serp serp H2O

O H2 O A hl 10 c H 2 rp

Pressure (GPa)

serp phase A

4

3

5.0 4.9

rp

O H2 A 10 serp

5

se

O H2 ”

6

“A

7

2O

phase A 10.6 10.1

O H2

O H2

700

800

900

1,000

1,100

Temperature (˚C)

Figure 6 Major phase stability boundaries in H2O-saturated peridotite and H2O contents stored in hydrous phases for harzburgite and lherzolite (upper and lower number, respectively; after Schmidt and Poli, 1998 and Fumagalli et al., 2001). The gray field denotes stability fields of hydrous phases. Serpentinized peridotite following the coldest P– T-path will not dehydrate up to transition zone depths, as phase boundaries of phase A (and “phase E” at .11 GPa, Frost, 1999) are flatter in P –T space than typical subduction P – T paths.

a significant increase of the bulk H2O content as most of the free fluid produced from serpentine breakdown would have escaped already. 3.17.5 DEVOLATILIZATION IN SEDIMENTS 3.17.5.1 Pelites Whereas subduction-zone metamorphism leads to a continuous decrease of H2O stored in MORB, this is not necessarily true for pelites at pressures up to 2.5– 3.0 GPa. By far the most abundant hydrous minerals in metapelites are the potassic micas phengite and biotite. The amount of H2O stored in micas (containing 9 –10 wt.% K2O and 4 – 4.5 wt.% H2O) is easily calculated from the bulk K2O content. Other important hydrous phases in subduction zone metapelites are talc, chloritoid, and chlorite (see compilation in Poli and Schmidt, 2002). The stabilities of these phases are strongly dependant on the XMg which varies widely in oceanic pelagic sediments, from 0.2 to 0.8 (e.g., Plank and Langmuir, 1998). Biotite is expected to transform completely to garnet þ phengite-bearing assemblages at pressures above

2.5 –3.0 GPa (Poli and Schmidt, 2002) but under high XMg biotite may persist to a maximum of 4 GPa (Hermann, 2002a). In contrast, phengite is stable to considerably higher pressures (up to 8 – 9 GPa, Domanik and Holloway, 1996; Ono, 1998). The absolute amount of hydrous phases, and therefore the maximum amount of H2O in metapelites, is also strongly controlled by the amount of quartz, which is highly variable in mica schists. An average of 9 metapelites (containing between 12 vol.% and 46 vol.% quartz) from the Nome Blueschist Terrane, Alaska (Thurston, 1985) yields 2.7 wt.% H2O stored in phengite, chlorite, paragonite, epidote, and glaucophane. The same average bulk composition would contain ,2.0 wt.% H2O stored in the hypothetical assemblage phengite—lawsonite—chloritoid (þ gar þ cpx þ coesite), which is stable near 3 GPa and 600 8C, and 1.1 wt.% H2O if phengite is the only hydrous phase at higher pressures. Thus, during subduction from 20 km to 80 km depth, 300 m of mica schist would yield only 2 –6% of the mass of H2O released from the underlying 2 km of MORB. These numbers are fairly approximate,

580

Generation of Mobile Components during Subduction

as our current knowledge of phase relationships in metapelites metamorphosed at high-pressure, lowtemperature conditions is still fragmentary, and further experimental work is required to explore the significance of talc, chlorite, and chloritoid in the critical pressure range between 2 GPa and 4 GPa. Nevertheless, metapelites contain generally less H2O than MORB in a lowpressure blueschist stage but more H2O at pressures beyond 2.5 GPa. Typically, only a few hundred meters of sediments may be subducted, and the quantity of volatiles stored in pelites is small when compared to MORB and serpentinized peridotite. The importance of pelites to the subduction factory lies in their relatively high concentration of K2O and other highly incompatible minor and trace elements, which may be concentrated in accessory phases (typically rutile, allanite, zircon, phosphates, ellenbergerite). Thus sediments can impart a strong trace-element signal to the slab fluids (or melts), which is quite distinct from that derived from the igneous oceanic crust. For example, if present, allanite (a phase that may be residual during melting) contains more than 90% of the whole-rock LREE and thorium; rutile contains more than 95% of titanium, niobium, and tantalum; zircon contains 95% of the wholerock zirconium and hafnium, and phengite, at a modal abundance of 20 –35% (Figure 8), incorporates more than 95% of the bulk-rock rubidium, barium, and caesium (see also Hermann, 2002b). Finally, as discussed above, pelites also impact melting, as carbonate-free pelites have the lowest melting temperature of all subducted lithologies (see Section 3.17.3.4) and, within the oceanic crust, are situated at the highest temperatures. 3.17.5.2 Carbonates Most of the discussion above centers on the breakdown of hydrous phases during subduction. The subducted mass of H2O in most subduction zones is much larger than that of CO2, and at many trenches, the sediments do not contain any significant carbonates. Near-trench sediment columns have molar CO2 : H2O ratios mostly below 1 : 3, and a ratio above 1 : 1 was only found along the middle-American to Peruvian margin (Rea and Ruff, 1996; Plank and Langmuir, 1998; compiled in Poli and Schmidt, 2002). If carbonates are confined to the sediments, they do not alter phase relations in the igneous part of the crust, as fluid flow is generally buoyancy driven away from the igneous crust. However, some carbonate precipitates during hydrothermal alteration (see Chapter 3.15) of the igneous crust. In these instances, carbonate is mainly found in veins and is much less pervasively distributed than H2O. For a more detailed understanding of fluid-producing

processes it is thus necessary to evaluate the effect of CO2 on subduction-related metamorphism and fluid transfer, on hydrous phase and carbonate stabilities, on fluid compositions, and how deep carbonates can subduct. In general, the succession of carbonate minerals with increasing pressure is calcite ! dolomite ! magnesite. XCO2 in the fluid decreases in this sequence. For low temperature P – T paths, aragonite replaces calcite. Both calculations (Kerrick and Connolly, 1998, 2001a,b) and experiments (Yaxley and Green, 1994; Molina and Poli, 2000) show that at low to intermediate temperatures (up to 800 8C at 4 GPa), a very small amount of CO2 saturates the system in carbonate and fluids are buffered to compositions with #10 mol.% CO2. Carbon dioxide becomes an important species in the fluid (to XCO2 values of 0.4 – 0.7) only at higher temperatures and comparatively low pressures in coexistence with calcite. Few experiments are available on the maximum pressure stability of carbonates in natural bulk compositions. Domanik and Holloway (2000) found magnesite in a calcareous pelite from 6.5 GPa to 11 GPa. In synthetic systems it is well known that carbonates are very stable at subsolidus conditions: between 5 GPa and 9 GPa dolomite breaks down along a curved reaction line (Luth, 2001) to magnesite and CaCO3 polymorphs, which are stable down to lower mantle pressures (Biellmann et al., 1993). Carbonates present within the subducting oceanic crust influence dehydration reactions, but contrary to what is intuitively believed, can enlarge the stability field of hydrous phases. This is especially true for melting reactions in mafic and pelitic compositions where a mixed H2O – CO2 fluid shifts melting reactions to higher temperatures and thus enlarges the amphibole and mica-stability fields. Carbonate minerals may remain stable in the presence of a siliceous melt and with increasing temperature decompose forming an immiscible carbonatite melt coexisting with a silicate melt. Molina and Poli (2000) demonstrated the effect of CO2 on phase relations in mafic compositions (Figure 3). With increasing pressure, the stable carbonates at 665– 730 8C are calcite (,1.4 GPa), dolomite, and dolomite þ magnesite (.2 GPa). Fluids coexisting with calcite are CO2 rich (XCO2 ¼ 0.4– 0.7) but at higher pressures, fluids coexisting with dolomite have much lower CO2 contents (XCO2 ¼ 0.02 – 0.2). This implies that carbon tends to fractionate into the solid with increasing pressure. The effect of CO2 on the stability of hydrous phases is surprising: in the carbonated system, plagioclase disappears at the same pressures as in the pure-H2O system (,1.5 GPa at 650 8C). However, its breakdown does not cause omphacite formation, which is

Devolatilization in Sediments delayed by ,0.5 GPa. Amphibole breaks down at ,2.5 – 2.6 GPa (Yaxley and Green, 1994; Crottini et al., 2003) and the fluid saturated solidus is located at ,730 8C at 2.2 GPa, as expected by the presence of H2O-rich fluids at such conditions. As XCO2 values are low (,0.15) in fluids produced by subduction zone metamorphism, the only efficient decarbonatization processes are either flushing with aqueous fluids from below or melting at relatively high temperatures. A scenario describing decarbonation reactions caused by fluid-infiltration during subduction would be a layer of carbonaceous sediment overlying hydrated oceanic lithosphere. At any time, the metacarbonate would probably contain a very small amount of equilibrated fluid that would have 5 – 15 mol.% CO2 at depths beyond 60 km (the exact CO2-fraction depending on pressure and temperature, Kerrick and Connolly (2001a,b) and Molina and Poli (2000)). Aqueous fluid produced in the serpentinite and basalt below the carbonate layer will migrate upwards, replacing the CO2bearing fluid, and then will locally equilibrate, i.e., consume some carbonates in order to increase the CO2 content in the fluid. If the aqueous fluid passes pervasively through the limestone, the entire carbonate layer will eventually be dissolved and the fluid migrating into the wedge will transport a significant amount of CO2 over time. For example, 100 m of average siliceous limestone (15.0 wt.% initial CO2; Plank and Langmuir (1998)) will be completely decarbonated from reaction with aqueous fluids derived from complete dehydration of 150 – 400 m of serpentinite (ignoring the basaltic crust) or 900 – 1,600 m of dehydrating MORB (ignoring the peridotite) at depths beyond 60 km, assuming pervasive infiltration and equilibration of the fluid. The entire CO2 thus derived will flow out of the carbonaceous sediment and into the mantle wedge. However, if the aqueous fluid is efficiently channeled, a very small fraction of the limestone would be affected (depending on channel-spacing) and .90% of the carbonates may survive complete dehydration of the underlying oceanic lithosphere. In this case, these carbonates will be subducted to depths of the transition zone.

3.17.5.3 Graywackes and Volcaniclastics A significant component of sedimentary columns at trenches are graywackes and volcaniclastic sediments of andesitic to dacitic composition (Plank and Langmuir, 1998). Phase diagrams of graywackes and andesites (Vielzeuf and Montel, 1994; Schmidt, 1993) suggest that, for our present purpose, these systems can be viewed as being intermediate between pelites and MORB. The significant K2O contents in graywackes and

581

volcaniclastics of intermediate composition lead to abundant micas (phengite or biotite), and thus reactions and phases similar to those in pelites. However, the CaO contents are significant enough for the formation of amphiboles, thus leading to phases and reactions similar to those in intrinsically CaO-rich MORB. Obviously, compared to MORB and pelites, reactions are shifted in P – T space due to different phase compositions, and the amounts of phases are highly variable.

3.17.5.4 Melting of Sediments Compared to Melting of MORB In hot subduction zones, melting of the downgoing oceanic crust might be achieved. In order to understand which lithologies might melt under what conditions it is necessary to compare melting relations of sediments and MORB under the various fluid-availability conditions. In this context, it should be noted that of all average sediments, a pelagic clay has the highest melt productivity. However, the initial melting temperature of most carbonate free sediments remains virtually identical as long as mica and quartz saturation is maintained. As for MORB, any carbonate addition to sediments results in a mixed H2O – CO2 fluid phase and an increase in melting temperatures by at least 100– 200 8C. The melting temperatures for MORB and pelitic sediments at water-saturated conditions are similar, the wet MORB solidus being situated typically 20 –50 8C above the wet pelite solidus (Figure 5). As pointed out above, fluid-saturated melting only becomes efficient (in terms of melt productivity) if fluid is added from an external source. In order to achieve flush-melting of the sediments and concomitant dehydration of the MORB layer, a very particular thermal field is necessary: first, the sediments have to pass 650 –800 8C at 1.5 –3.0 GPa; secondly, the temperature within the lower part of the MORB layer must be 100– 200 8C lower than in the sediments in order to provide a significant quantity of fluid through chlorite- and amphibole-decomposition reactions. Within the temperature range where flush melting would occur in the sediments, and considering a normal temperature distribution (i.e., temperature decreasing with depth in the oceanic crust), serpentinized peridotite would remain in the chlorite þ serpentine stability field and thus not produce any fluid. For the same reason, flush melting of MORB near the wet solidus by fluids derived from underlying serpentinite is highly unlikely in the depth range of interest, as temperatures in the serpentinized peridotite should remain 50 – 200 8C lower than in the MORB (compare to Figure 1). Thus, conditions necessary for flush melting might be

582

Generation of Mobile Components during Subduction

realized in a particular subduction zone; however, flush melting cannot be treated as the general case and is not expected for the fast subduction zones dominant in the Pacific rim. For these, thermal models predict temperatures far too low for any melting in the 1.5– 3.0 GPa depth range (Davies and Stevenson, 1992; Kincaid and Sacks, 1997; Gerya et al., 2002). Fluid-absent melting involving biotite in sediments or amphibole in MORB produces melts with distinct compositions at different P – T conditions (Figure 5), with different pressure dependencies of the melting reactions. The amphibole-melting reaction bends strongly back around 2 GPa and near the solidus, amphibole disappears at 2.4 GPa (Figure 5). In MORB, epidote/zoisite remains stable at the solidus to ,3 GPa, leaving a possibility for minor fluidabsent melting up to these pressures. Any melting at higher pressures depends on the K2O content of MORB and the related presence of minor phengite. In average metapelites and graywackes, biotite is completely replaced by phengite at 2.5– 2.8 GPa (Auzanneau, unpublished experiments, see also Vielzeuf and Schmidt (2001)) but at high Mg/(Mgþ Fe) biotite stability extends to ,4 GPa (Hermann, 2002a). An interesting feature of melting at pressures above 3 GPa arises from the fact that MORB, pelites, and also intermediate andesite or graywacke compositions all contain garnet, cpx, phengite, and coesite (Schmidt, 1996; Okamoto and Maruyama, 1999; Hermann and Green, 2001; Schmidt and Vielzeuf, 2001). Peraluminous graywackes and pelites have kyanite in addition. Thus, all the lithologies of the oceanic crust contain the same assemblage and fluid-saturated melting occurs through the identical reaction: phengite þ coesite þ cpx þ H2O ¼ melt. If a significant amount of free water is not available, melting occurs through phengite þ cpx ¼ garnet þ melt ^ kyanite. This reaction takes place ,150– 200 8C above the wet solidus (at 3 GPa to at least 5 GPa) and leads to 20 –30 wt.% melt in the metasediments and to a few percent melt in the mafic rocks (dependent on the bulk K2O content). In MORB, phengite is immediately consumed upon melting and the temperature must rise by .100 8C to significantly increase melt fractions through the reaction: cpx ¼ garnet þ melt (25 – 30% melt). Nevertheless, temperatures necessary for such fluid-absent, high-pressure melting are high and unrealistic in most subduction zones. The most likely melting scenarios for subducted oceanic lithosphere strongly depend on the thermal gradient within the subducted crust. In the depth range of interest, most thermal models (e.g., Davies and Stevenson, 1992; Furukawa, 1993; Kincaid and Sacks, 1997;

Gerya et al., 2002) predict a temperature difference of #200 8C from top of the sediments to the bottom of the crust (which at the same time is the top of the serpentinized peridotite layer). With such a gradient, significant fluidsaturated sediment flush melting cannot take place due to fluids originating from the directly adjacent MORB layer. However, the temperatures necessary to obtain fluids from the serpentinized peridotite cause the MORB layer (intercalated between sediments and serpentinite) to be at temperatures above the fluid-saturated solidus. Thus it appears likely to achieve either no significant fluid-saturated melting at all, or flush melting from both sediments and MORB. For depths of 50– 80 km, achieving the temperatures necessary for fluid-absent melting of sediments and MORB results in a temperature that causes dehydration in the serpentinized peridotite layer. This again would enhance melt productivities through flush melting.

3.17.6 SERPENTINIZED PERIDOTITE The serpentinized peridotite layer situated just below the igneous oceanic crust (or often brought to the surface in slow spreading oceans like the Atlantic) constitutes a major H2O reservoir in subducted lithosphere, of comparative size to the oceanic crust. It is difficult to estimate the amount of serpentinization within this layer. The only certainty is that the degree of serpentinization is highly variable both on a regional and local scale. Hydrothermal systems near the ridge and transform faults of all dimensions are the primary sites of hydration, which is mostly serpentinization. Recently, hydration has also been suggested to occur in extensional faults that run parallel to the trench and are caused by bending of the subducting plate at the onset of subduction (Peacock, 2001). Our best estimate gives an average of 20% serpentinization to a few kilometer depth (Schmidt and Poli, 1998). Although there might be some localized serpentinization along faults much deeper in the oceanic lithosphere, largescale serpentinization and the resulting lowdensity peridotite (2.3 –2.5 g cm – 3) would cause a buoyancy problem during subduction. Whereas a number of reactions at intermediate to elevated temperatures are important for hydration of peridotite in the overlying mantle wedge directly adjacent to the top of the oceanic crust (Figures 6 and 7), almost any subduction P – T-path will keep the slab in the serpentine stability field to at least .2 GPa (Ulmer and Trommsdorff, 1995). As a consequence, hydrated peridotite in the downgoing lithosphere will remain as serpentine and chlorite (þ olivine þ clinopyroxene) while a multitude of reactions is

Implications for Trace Elements and an Integrated View of the Oceanic Lithosphere taking place in the oceanic crust above. Thus, the serpentinized peridotite layer of the oceanic lithosphere does not produce any significant fluid up to pressures of 3 –6 GPa, where, depending on temperature, serpentine breakdown may occur (Figure 6). The H2O contents of Figure 6 are calculated for average harzburgite. However, oceanic alteration does not only add H2O, but also removes MgO. Peridotite sitting just above the slab surface experiences alteration by fluids derived from generally quartz-saturated sediments and MORB and thus is likely to become somewhat SiO2 enriched (Manning, 1994). If the MgO/ (MgO þ SiO2) ratio in the peridotite is shifted from lying between olivine and serpentine to a value between serpentine and talc, then talc becomes stable to higher pressures (talc þ serpentine reaction; Ulmer and Trommsdorff (1999)). In monomineralic veins, talc might persist to its maximum-pressure stability limit at 4.5 – 5.0 GPa (Pawley and Wood, 1995). Subduction zones can be divided into two types: those where the serpentine-out reaction is crossed below 6 GPa (and major dehydration of serpentinized peridotite occurs), and those where the serpentine stability boundary is crossed at .6 GPa. As discussed above, almost no fluid production is expected when serpentine reacts to phase A (see Figures 6 and 7) and the H2O stored in the hydrated peridotite is expected to subduct deeply. In contrast, if the stabilities of serpentine and phase A do not overlap along a given slab geotherm, a significant flushing zone (over 20 – 30 km depth) is expected at the transition between serpentine-chlorite peridotite ! chlorite peridotite ! anhydrous garnet peridotite. The width of such a zone is controlled by the actual stability of chlorite, which extends ,100 8C higher than the stability of serpentine in complex systems approaching natural rocks (Fumagalli, 2001). This picture is somewhat modified when the P – T-path produces the sequence serpentine ! ˚ phase ! phase A. In this case, a moderate 10 A amount (max. 0.8 wt.%) of H2O subducts to great depth, while .75% of the initially subducted H2O of the serpentinized peridotite is lost via dehydration. In a number of arcs (NE Japan – Kuriles – Kamchatka, Aleutians, N. Chile), so-called double seismic zones are observed. Whereas the upper seismic zone correlates with the oceanic crust, it has been suggested (Seno and Yamanaka, 1996; Peacock, 2001) that the lower seismic zone corresponds to the limit of serpentine stability in the lower part of the oceanic lithosphere (see Figure 7). It has been argued that the lower seismic zone earthquakes are triggered by reactivation of ancient faults through fluid saturation, where the fluids derive from serpentine dehydration.

583

3.17.7 IMPLICATIONS FOR TRACE ELEMENTS AND AN INTEGRATED VIEW OF THE OCEANIC LITHOSPHERE Trace elements and isotopes such as 10Be, B, Li, Ba, Cs, and elements of the U –Th-series provide important information on main element-transfer processes in arcs and are reviewed by Morris (see Chapter 2.11). Such studies allow one to distinguish between a sediment and an altered MORB signal in island arc volcanics (Morris et al., 1990), to deduce across-arc variations as a function of slab depth (Ishikawa and Nakamura, 1994; Ryan et al., 1995; Moriguti and Nakamura, 1998), yield time constraints on element transfer (Sigmarsson et al., 1990; see Turner et al., 2003, for review), and allow one to distinguish between dehydration and melting processes (Sigmarsson et al., 1998; Martin, 1999). However, the correct interpretation of their concentrations and spatial distributions rely on major element phase relations and on the understanding of mobile phase production discussed in this chapter.

3.17.7.1 Mobile Phase Production and Trace-element Transfer Dehydration, or more generally, devolatilization of the oceanic crust is a process that combines continuous and discontinuous reactions in a variety of heterogeneous bulk compositions. In addition, within a vertical column—the sedimentary, mafic, and serpentinized peridotite layers— each experience a significant thermal gradient. The result is a continuous, but not constant, production of a fluid or melt, with the rate of mobile phase production generally decreasing with depth. Peaks in the volatile flux result from significant discontinuous reactions. However, despite the continuous fluid flux, trace elements may not necessarily be released continuously. All dehydration reactions in oceanic lithosphere take place at temperatures where diffusion rates in most minerals are insignificant compared to the available time span for fluid production in subduction zones. In terms of trace-element partitioning it is thus necessary to distinguish between the mineral mode in a given composition and the reactive volume, which will be much smaller. For example: a MORB at 4 GPa, 700 8C has 48 vol.% garnet, 39 vol.% cpx, 5 vol.% lawsonite, 2 vol.% phengite, and 6 vol.% coesite. The breakdown of lawsonite can be modeled via the reaction lawsonite þ clinopyroxene þ garnet1 ¼ garnet2 (grossular enriched) þ H2O, which produces ,8 vol.% garnet in the rock (Schmidt and Poli, 1998). The garnet that grows from the breakdown of lawsonite is in equilibrium with the fluid. However, diffusion rates at 700 8C

584

Generation of Mobile Components during Subduction B+G: P:

Subducted mass of water in Oceanic lithosphere at 20 km depth:

4.9–5.9 108 g m–2 1.8–3.2 108 g m–2 80–150 km above slab surface

0

200 400 600

serp 3.7–4.7

amp chl

50

800

am p

AMP

37–47%

1,000

ser

80 km

p

Depth (km)

20 km

1.4–1.6

1,200

100

18–21%

cld

chl law

150

150 km

1.6–2.8

21–35%

ser

p

200

4 cm a–1 30˚

10

A

phe Dehydration rates (106 (g m–2) km–1)

2.2

7.0

3.0

220 km 0%

(a) 80–150 km above slab surface 0

serp

1.6–2.9

200 400 600

chl amp

800

50

20 km 21–32%

80 km

ser

100

1.8–3.1

p

1,200 20–40%

cld

l

ch

Depth (km)

p

am

1,000

150 km de

150

No ation dr hy

4–9% p

ser

(b)

8 cm 60˚

a–1

A 10

200

0.4–0.6

“A” law + phe

3.7

3.4

–200 km 2.5–3.9

32–43%

0.7

Figure 7 Stabilities of hydrous phases and masses of H2O in a unit section (in 108 g m – 2) for two typical temperature fields during intra-oceanic subduction (a) for an intermediate burying velocity (4 cm yr – 1 convergence rate and 308 dip angle of the slab) and (b) a fast and cold subduction zone (8 cm yr – 1 convergence rate and 608 dip angle; isotherms from Furukawa, 1993). Dark gray: oceanic crust, light gray: lithosphere of the overriding plate in the Furukawa (1993) model. The dotted lines denote maximum stability fields and not effectively hydrated peridotite. In a subduction zone with a fast subduction rate, serpentine and phase A fields overlap, preventing significant dehydration

Implications for Trace Elements and an Integrated View of the Oceanic Lithosphere for garnet and cpx are so slow that in the available time span (several tens of thousands of years, based on radioactive disequilibrium, see review by Turner et al., 2003) the volumes of the nonreacting garnet and cpx affected by diffusion are negligible (,0.1 vol.%). Thus, the trace elements formerly residing in lawsonite and the 3 vol.% of cpx that decomposed with lawsonite, will be redistributed between the newly formed garnet and fluid. However, .85 vol.% of the rock is not equilibrated and thus should not be included when calculating geochemical residua or the traceelement content of the fluid. As a consequence of slow solid-state diffusion, in most cases the only elements that may be mobilized are those that are hosted in minerals that decompose. It is thus necessary to establish trace-element residence in subduction-zone lithologies (Figure 8). The following cases may be distinguished: .

.

.

.

.

A given trace element, for example, boron (Ryan et al., 1995), is extremely soluble in the fluid and any mineral/fluid-partition coefficient is p1. Such elements may show a continuously decreasing concentration in the fluid with increasing depth. Concentrations might already be very low in fluids produced at moderate depths due to a shallow effective removal from the subducting crust. A given trace element has partition coefficients close to unity and thus its concentration doesn’t change in the fluid as subduction progresses. A given trace element is strongly partitioned into a particular hydrous phase (e.g., cerium, strontium into lawsonite or epidote; Figure 8) and has only moderate-to-low cpx/fluid and garnet/fluid partition coefficients. These elements will quantitatively enter into the fluid at the breakdown reaction(s) of the given mineral and cause a variation in their concentrations that is not at all proportional to the fluid-flux. A given trace element (e.g., barium, and to some extent beryllium in the sediments) partitions into mica and as mica is dissolved away with increasing depth and temperature, enters into the fluid only at greater depth, its flux at low temperature/pressure being small. A given trace element strongly partitions into cpx/garnet and will be returned to replenish the deep-mantle trace-element reservoir.

585

3.17.7.2 Integrating Fluid Flux over the Entire Subducted Oceanic Crust: An Example The layered structure of oceanic lithosphere originating from fast-spreading ridges, combined with temperature gradients within the subducting lithosphere define the relationship between the amount and depth range of fluids and/or melts generated in the sedimentary and mafic layers and fluids generated from the ultramafic layer. The interdependence between sediment and

Figure 8 Distribution of Be, B, Rb, Sr, Y, Ce, and Ba between minerals of average MORB and pelite in blueschist and eclogite facies (employing representative mineral modes for natural blueschists and experimental epidote-eclogite, and trace element concentration data mostly from Domanik et al., 1993). At subsolidus temperatures, diffusive equilibration is ineffective (except for micas) and the equilibrating volume that needs to be taken into account for trace-element modeling is defined by the reacting minerals. Thus, a given trace element equilibrates with the fluid only when its host phase(s) break(s) down.

of the serpentinized peridotite of the oceanic lithosphere. In such a thermal situation, a thin layer of hydrated peridotite is stable above the oceanic crust and drags some H2O downwards. Details of the calculations: the input is set to H2O-amounts at 20 km subduction depth (upper left values), the oceanic lithosphere is composed of 2 km of fully hydrated basalt, 5 km of gabbro hydrated to 20 vol.% and 5 km of serpentinized peridotite with 10 – 20 vol.% hydrated. Labels on open arrows and percentages at the right side are masses of H2O (fluxing the mantle wedge) integrated over the depth ranges 20– 80 km, 80– 150 km, and 150 – 220 km, and the mass (108 g m – 2 of subducted crust) of H2O subducting beyond 220 km. The arrows below (a) and (b) give dehydration rates in 106 g m – 2 per (vertical) kilometer of subduction at 50 km, 100 km, and 180 km depth. The highest uncertainty in terms of H2O masses is in the estimate of initially subducted H2O in the peridotite of the oceanic lithosphere.

586

Generation of Mobile Components during Subduction

MORB melting and dehydration in serpentinized peridotite was illustrated in Section 3.17.5.4. Here, an example of the effects on trace-element transfer as a function of fluid : rock ratios is illustrated. The layered structure of the oceanic lithosphere may cause the fluid : rock ratio in the sedimentary layer to be greater than one (as fluids from the underlying mafic and peridotitic layers must rise through the sediments). Thus, some of the trace elements (e.g., beryllium, thorium) commonly considered to be only efficiently mobilized by melts could also be quite effectively mobilized by fluids, if the entire subducted lithosphere is considered. Recently, an apparent contradiction was put forward to argue for melting of sediments contemporaneously with dehydration of MORB. It was estimated that .30– 40% of the subducted beryllium and thorium, which are strongly enriched in sediments, are recycled into the mantle and extracted to the surface via arc volcanism (Johnson and Plank, 1999, and references therein). At the same time, boron, which is strongly enriched in altered MORB, and uranium appear to be effectively recycled into arc magmas by fluids. Based on bulk partition coefficients xtl/fluidD of 2 –4.8 and xtl/meltD of 0.7– 1.5 for beryllium and thorium, Johnson and Plank (1999) argued that recycling of .30 –40% of these elements is only possible when sediments melt while MORB dehydrates (mobilizing boron and uranium). It was then shown by thermal modeling that such a temperature distribution is possible (Van Keken et al., 2002). However, this scenario considers sediments, MORB, and hydrated peridotite as independent systems. Considering that the bulk crystal-fluid partition coefficients for beryllium and thorium are not extremely high, i.e., 2 –5, and considering that most of the beryllium partitions into phengite, and that phengite is the only phase where diffusive reequilibration at sub-solidus temperatures is possible, it follows that an elevated fluid/rock ratio can dissolve .80% of the beryllium (and probably thorium) in the sediments, if pervasive fluid infiltration from the lower layers occurs. Employing the above partition coefficients, the fluid produced from an average thickness MORB layer through dehydration below 3 GPa would be sufficient for leaching 60% of beryllium (and possibly thorium) out of 200 m of pelagic sediments. Thus, sufficient beryllium could be mobilized before typical melting depths are reached. This estimate is conservative because partition coefficients for relatively low-pressure fluids are employed. Solute-rich, high-pressure fluids and continuous dissolution of phengite would greatly facilitate the transfer of beryllium to the mantle wedge. The trace-element transport capacities of such solute-rich fluids are expected to increase

significantly compared to low-pressure, lowdensity fluids; however, the little experimental data that are available are not enough to suggest this. In fact, significant mobilization of beryllium could simply be taken as evidence that micas dissolve away in a dissolution regime. 3.17.8 CONCLUSIONS AND OUTLOOK The complexity of natural processes—i.e., heterogeneous bulk compositions, heterogeneous volatile distribution, equilibrium on different scales, and kinetic effects—may fairly complicate individual subduction zones. Nevertheless, there is no reason for pessimism. The dehydration behavior of the two volumetrically predominant lithologies, basalt and peridotite, are fairly predictable. The resulting mobile component will then have a sediment signal added. This signal can be reasonably well defined when the subducting column of sediments is well known; there is ample geochemical evidence that the efflux in arc magmas is fairly proportional to the influx in terms of subducted sediment component (e.g., Plank and Langmuir, 1993; Morris, 2003; Turner et al., 2003). The complexity of natural processes in subducting slabs is reflected by the complex distribution of volcanic emissions in subduction zones. First, it should be emphasized that 69% of modern subduction zones on the Earth (40,900 km, unpublished compilation) show active volcanism in the Quaternary. The rest (18,500 km) are not volcanic because of either unfavorable thermomechanical environments (e.g., flat slabs, initiation of subduction, etc.) or possibly the lack of a sufficient amount of volatiles released at depths where melting could take place. Even though we believe that there is no straightforward relationship between the location of fluid release in the slab and volcanic emissions, the variability of reaction patterns illustrated in this chapter is recorded by variability in the distribution of volcanic arcs on the Earth’s surface (Figure 9). The spatial onset of volcanism, the arc, is probably primarily controlled by the thermal structure in the mantle wedge, i.e., when a sufficient thickness and convection intensity in the mantle wedge is reached in order to allow temperatures necessary for the formation of primitive arc magmas (Kushiro, 1987; Schmidt and Poli, 1998). The depth of the slab below the volcanic front is often regarded as a relevant parameter to characterize petrologic processes occurring in subduction zones (Gill, 1981; Tatsumi, 1986). This parameter has often been expressed as single-valued with some sort of standard deviation (128 ^ 38 km Gill, 1981; 110 ^ 38 km Tatsumi, 1986; 108 ^ 14 km for the trench side of the chain, 173 ^ 12 km for the back-arc side of the

Conclusions and Outlook chain, Tatsumi and Eggins, 1995) on the basis of a fairly arbitrary volcano counting in selected arcs. On the contrary, our compilation is based on recent geophysical acquisitions (slab surface tomography) and on the spatial extent of volcanic activity and shows that, although some maximum is found at ,100 km depth, a continuum in “depth of the slab surface below the volcanic front” is observed (Figure 9(a)). Moreover, despite a majority of volcanic arcs being fairly well focused 10,000

587

on the surface (less than 50 km wide), volcanic activity over more than 100 km arc width is not unusual (Figure 9(b)). Such an arc width corresponds to a comparative depth range of the slab surface. In some continental arcs (e.g., Sumatra, the Southern Volcanic Zone in the Andes), a narrow volcanic arc is the result of a position of virtually all volcanoes on a major fault subparallel to the trench. In such a case, it appears likely that the volcano distribution is not directly related to 8,000

8,000 Length (km)

6,000 6,000 4,000 4,000 2,000 2,000

0 40 60 80 100 120 140 160 180 200 220 240 260 Depth of slab surface below volcanic front (km) (a)

0 0 (b)

40

80

120 160 200 240 280 320 360 Width of volcanic arc (km)

Figure 9 Subduction zone statistics: histograms of depths of (a) the slab surface below the volcanic front and (b) the width of volcanic arcs. The vertical axis denotes arc lengths in km measured at the trench. This is our own compilation (unpublished) based on locations of quaternary volcanoes and slab surfaces from tomography (if available), otherwise earthquake depths. Volcanic arc Backarc

Forearc

0 Wherlites, pyroxenites

Pressure (GPa)

75

1,400 ºC Partially molten Flow lines region

3 4

100

5 6

Chlorite Phlogopite

200

10Å

8

“A”

9

Lawsonite Phengite

150

Antigorite

K-richterite

7

10

talc 50

1,100 ºC

2

Figure 10

amp

0

Depth (km)

1

Serpentine diapirs

250

Devolatilization

300

Cartoon of events in a typical cold subduction zone in which melting of the oceanic crust does not occur (after Poli and Schmidt, 2002).

588

Generation of Mobile Components during Subduction

processes taking place below the lithosphere of the overriding plate. It is beyond the scope of this review to discuss in detail the statistical parameters of subduction zones, and Figure 10 is intended to demonstrate that conditions attained in subducting slabs are highly variable, even for similar convergence parameters, and that interplay between thermomechanical properties and reaction paths are responsible for a complex pattern of fluid release and magma genesis. Furthermore, any correlation of kinematic subduction parameters with volcano location tacitly assumes steady state, which is not necessarily the case. At present, the resolution of thermal models for the oceanic crust with its large temperature gradient is not sufficient. This is because the thermal field strongly depends on the flow field (and on the degree of mechanical coupling between the subducting slab and the down-dragged mantle wedge), which in turn depends on the viscosity, which is a function of P and T and the materials present. The latter then depend on the chemical reactions taking place, which in turn depend on the temperature field. A temperature distribution model is thus a fairly complex problem and needs input from a large variety of disciplines. An example of feedback between temperature field and phase petrology is serpentinization in the mantle wedge directly overlying the subducted slab (Gerya et al., 2002). First, the calculated temperature distribution from thermomechanical models results in a pressure – temperature region where serpentine would be potentially stable. Second, dehydration rates and fluid transport mechanisms (pervasive versus channeled flow) allow one to model the amount of serpentine formed. However, the serpentinized peridotite has four to six orders of magnitude lower viscosity than dry peridotite. This strongly influences the coupling of the downgoing slab and the convection of the mantle wedge. It also influences the possible amount of shear heating at the slab surface. If convection patterns change due to a modified rheology caused by phase transformations, the temperature field changes and the permissible region for serpentine changes. We thus need more complex models in which these parameters are varied simultaneously. Upwards directed interaction between the different lithologies largely depends on the temperature gradient within the oceanic lithosphere, which is among the most difficult parameters to model. Thus, at present, a purely forward model for mass transfer in a given subduction zone is not reliable, geochemical information on the subduction output (see Chapters 2.11 and 3.18) is necessary to constrain likely mass-transfer processes.

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q 2003, Elsevier Ltd. All rights reserved No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without prior written permission of the Publisher.

Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 567–591

3.18 One View of the Geochemistry of Subduction-related Magmatic Arcs, with an Emphasis on Primitive Andesite and Lower Crust P. B. Kelemen and K. Hanghøj Woods Hole Oceanographic Institution, MA, USA and A. R. Greene Western Washington University, Bellingham, WA, USA 3.18.1 INTRODUCTION 3.18.1.1 Definition of Terms Used in this Chapter 3.18.2 ARC LAVA COMPILATION 3.18.3 CHARACTERISTICS OF ARC MAGMAS 3.18.3.1 Comparison with MORBs 3.18.3.1.1 Major elements 3.18.3.1.2 We are cautious about fractionation correction of major elements 3.18.3.1.3 Distinctive, primitive andesites 3.18.3.1.4 Major elements in calc-alkaline batholiths 3.18.3.2 Major and Trace-element Characteristics of Primitive Arc Magmas 3.18.3.2.1 Primitive basalts predominate 3.18.3.2.2 Are some low Mg# basalts primary melts? perhaps not 3.18.3.2.3 Boninites, briefly 3.18.3.2.4 Primitive andesites: a select group 3.18.3.2.5 Three recipes for primitive andesite 3.18.3.3 Trace Elements, Isotopes, and Source Components in Primitive Magmas 3.18.3.3.1 Incompatible trace-element enrichment 3.18.3.3.2 Tantalum and niobium depletion 3.18.3.3.3 U-series isotopes 3.18.3.3.4 Geodynamic considerations 3.18.4 ARC LOWER CRUST 3.18.4.1 Talkeetna Arc Section 3.18.4.1.1 Geochemical data from the Talkeetna section 3.18.4.1.2 Composition, fractionation, and primary melts in the Talkeetna section 3.18.4.2 Missing Primitive Cumulates: Due to Delamination 3.18.4.3 Garnet Diorites and Tonalites: Igneous Garnet in the Lower Crust 3.18.5 IMPLICATIONS FOR CONTINENTAL GENESIS 3.18.5.1 Role of Lower Crustal Delamination in Continental Genesis 3.18.5.2 Additional Processes are Required 3.18.5.2.1 Andesitic arc crust at some times and places 3.18.5.2.2 Arc thickening, intracrustal differentiation, and mixing

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3.18.6 CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

3.18.1 INTRODUCTION This chapter has four main aims. (i) We wish to provide a comprehensive picture of the composition of volcanic rocks from subduction-related magmatic arcs. There are several recent reviews of the geochemistry of arc basalts. This chapter differs in including andesites as well as basalts, in focusing on major elements as well as trace elements and isotopes, and in using elemental abundance in “primitive” lavas, rather than traceelement ratios, to investigate enrichments of incompatible elements in arc magmas relative to primitive mid-ocean ridge basalts (MORBs). (ii) We review evidence in favor of the existence of andesitic as well as basaltic primary magmas in arcs. While we have recently reviewed evidence for this in the Aleutian arc, in this chapter we broaden our data set to arcs worldwide, and concentrate on whether mixing of lower crustal melts with primitive basalts offers a viable alternative to the hypothesis that there are “primary” andesites, i.e., andesites in Fe/Mg equilibrium with mantle olivine, passing from the mantle into the crust beneath arcs. (iii) We present new data on the composition of arc lower crust, based mainly on our ongoing work on the Talkeetna arc section in south central Alaska. To our knowledge, this is the first complete ICP-MS data set on an arc crustal section extending from the residual mantle to the top of the volcanic section. (iv) We summarize evidence from arc lower crustal sections that a substantial proportion of the dense, lower crustal pyroxenites and garnet granulites produced by crystal fractionation are missing. These lithologies may have been removed by diapirs descending into less dense upper mantle. In order to achieve these aims, and to limit the length of the chapter, we have not provided a detailed review of theories regarding the origin of primitive arc basalts, the mixing of magmatic components derived from the upper mantle, aqueous fluids, and sediment melts, or opensystem processes in the crust including mixing and assimilation. For a more complete view of these theories, we refer the reader to the many excellent review papers and individual studies that are, all too briefly, cited below. 3.18.1.1 Definition of Terms Used in this Chapter The following terms are used extensively throughout this chapter, and/or in the recent arc

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literature. We have tried to define each term where it first arises in the text. However, we realize that not all readers will read every section. Thus, we include the following brief definitions here. Accumulated minerals: These are crystals in lavas that are in excess of the amounts that could crystallize from a melt in a closed system. Like cumulates, some cases of accumulated crystals may be impossible to detect, but others stand out because they result in rock compositions different from all or most terrestrial melts. Adakite: This term is justly popular, but unfortunately it means many different things to different people, so we try not to use it. It generally is used for andesites and dacites with extreme light rare earth element (REE) enrichment (e.g., La/Yb . 9), very high Sr/Y ratios (e.g., Sr/Y . 50), and low yttrium and heavy REE concentrations (e.g., Y , 20 ppm, Yb , 2 ppm). Andesite: For the purposes of this chapter, andesites are simply lavas (or bulk compositions) with .54 wt.% SiO2. For brevity, we have not subdivided relatively SiO2-rich magmas into dacite, rhyodacite, etc. Basalt: In this chapter, basalt means lavas or bulk compositions with ,54 wt.% SiO2. Note that we have eliminated alkaline lavas (nepheline or kalsilite normative) from the data compilation. In a few places, we use terms picrite and komatiite to refer to basaltic melts with more than 15 wt.% and 18 wt.% MgO, respectively. Boninite: By boninites, we mean andesites with TiO2 , 0.5 wt.%, plus lavas identified as boninites in the original papers presenting geochemical data (includes some basalts!). Calc-alkaline: Magmas having both high Na þ K at high Mg# (Irvine and Baragar, 1971) and high SiO2 at high Mg# (i.e., high SiO2 at low Fe/Mg, Miyashiro (1974)); “tholeiitic magmas” have lower SiO2 and Na þ K at the same Mg#, when compared to calc-alkaline magmas. Compatible element: An element with equilibrium solid/melt partition coefficient .1. In garnet (and zircon), heavy REE, scandium, vanadium, and yttrium are compatible. These elements are incompatible in all other rock-forming minerals involved in igneous fractionation processes discussed in this chapter. Cumulate: A rock formed by partial crystallization of a melt, after which the remaining melt was removed. While some cumulates may be difficult to recognize, others are evident because the cumulate mineral assemblage has major and/or trace-element contents distinct from all or most

Arc Lava Compilation terrestrial melts, easily understood as the result of crystal/melt partitioning. Eclogite facies: This term represents high-pressure and relatively low temperature, metamorphic parageneses with omphacitic clinopyroxene and pyrope-rich garnet. EPR: East Pacific Rise. Evolved: Lavas, melts, liquids with Mg# , 50. JDF: Juan de Fuca Ridge. High Mg#: Lavas, melts, liquids with Mg# from 50 to 60. Incompatible element: Equilibrium solid/melt partition coefficient ,1. LILEs: Large ion lithophile elements, rubidium, radium, barium, potassium (and caesium— but we do not use caesium data in this chapter). Although they are not, strictly speaking, LILE, we sometimes group thorium and uranium with the LILE when referring to elements that are highly incompatible. Mantle wedge: “Triangular” region underlying arc crust, overlying a subduction zone, extending to perhaps 400 km depth. Mg#: 100 £ molar MgO/(MgO þ FeO), where all iron is treated as FeO. MORB: Mid-ocean ridge basalt. Primary: Lavas, melts, and liquids derived solely via melting of a specific, homogeneous source. In practice, it is hard to recognize or even conceive of a truly primary melt. Strictly speaking, even mantle-derived MORBs may be mixtures of primary melts derived from a variety of sources, including polybaric melts of variably depleted peridotites and/or “basaltic veins.” For brevity we have used the term “primary” in a few cases in this chapter. Where it is used without qualification, we refer to melts that are, or could be, in equilibrium with mantle olivine with Mg# ¼ 90– 93. Primitive: Lavas, melts, and liquids with Mg# . 60. Primitive cumulates have Mg# . 85. REE groups: Light—lanthanum to samarium; middle—europium to dysprosium; heavy— holmium to lutetium. Subduction-related magmatic arc (or simply, arc): Chains of volcanoes on the overthrust plate parallel to, and ,100 –200 km horizontally away from, the surface expression of a subduction zone, together with coeval, underlying plutonic rocks.

3.18.2 ARC LAVA COMPILATION Data in this and subsequent sections are from the Georoc database at http://georoc.mpch-mainz. gwdg.de/ (arcs worldwide), our Aleutian arc compilation, including all data available from the database compiled by James Myers and Travis McElfrish and available at http://www.gg.uwyo. edu/aleutians/index.htm, supplemented by

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additional data cited in Kelemen et al. (2003b), a new Costa Rica and Panama compilation (Abratis and Worner, 2001; Carr et al., 1990; Cigolini et al., 1992; de Boer et al., 1991, 1988, 1995; Defant et al., 1991a,b, 1992; Drummond et al., 1995; Hauff et al., 2000; Herrstrom et al., 1995; Patino et al., 2000; Reagan and Gill, 1989; Tomascak et al., 2000), and Central American data compiled by Mike Carr and available at http://www.rci.rutgers.edu/~carr/index.html. We also included a very complete data set on lavas from Mt. Shasta in the southern Cascades (Baker et al., 1994; Grove et al., 2001). We compare compositions of arc lavas to data on MORB glasses downloaded from PetDB, online at http://petdb.ldeo.columbia.edu/petdb/. Lava data come from intra-oceanic arcs (Tonga, n ¼ 704; Kermadec, 189; Bismark/New Britain, 165; New Hebrides, 252; Marianas, 834; Izu-Bonin, 878; oceanic Aleutians, 1082; South Sandwich, 328; Lesser Antilles, 356) and arcs which are, or may be, emplaced within older continental material or thick sequences of continentally derived sediment (Philippines, 221; Indonesia, 380; Papua New Guinea, 78; SW Japan, 92; NW Japan, 2314; Kuriles, 721; Kamchatka, 447; Cascades, 202; Central America, 857; the Andes, 1156; Greater Antilles, 175). Notably missing are data from New Zealand, the Alaska Peninsula, and Mexico. We apologize to authors whose work is not cited here, but whose analyses we compiled using large, online databases. It is simply not practical to cite the sources of all the data compiled for this chapter. We urge readers to contact us, and to visit the online databases, in order to check on the provenance of specific data. In comparing arc lavas to MORB glasses, it is important to keep in mind that none of our arc data sets discriminate between true liquid compositions, and compositions of lavas potentially including abundant, accumulated phenocrysts. Many of the lavas in our complete compilation had MgO contents .20 wt.% at 100 MgO/ (MgO þ FeOt), or Mg#, . 65. Samples with more than 20 wt.% MgO have been eliminated from all of our plots; we believe these, and possibly many other lavas with 10 –20 wt.% MgO, reflect the effect of accumulated olivine. In addition, many high Mg# lavas with lower MgO contents contain abundant phenocrysts of clinopyroxene and/or olivine. The high Mg#s of such samples could be due to accumulated clinopyroxene or olivine, at least in part. High Mg# andesites play a large role in the interpretive sections of this chapter, and thus this problem should be kept in mind. With this said, the high SiO2 and alkali contents, and the low MgO and CaO contents, of these samples, and the similarity of nearly aphyric primitive andesites from Mt. Shasta

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(Baker et al., 1994; Grove et al., 2001) to other, similar compositions worldwide, does not allow for much accumulated olivine and clinopyroxene. Finally, abundant accumulated plagioclase probably accounts for some lava compositions with very high Al2O3 contents (e.g., Brophy, 1989; Crawford et al., 1987). However, because this may remain controversial, we did not eliminate lavas on the basis of Al2O3 content. Another issue is that when relying on compiled information from databases, one has to be aware of the possibility that data may have been incorrectly entered, transferred, or normalized. Thus, for visual clarity, we eliminated some outliers, including lavas with more than 80 wt.% SiO2 and Mg# . 50, lavas with more than 70 wt.% SiO2 and Mg# . 60, lavas with less than 10 wt.% MgO and Mg# . 80, and lavas with less than 5 wt.% MgO and Mg# . 75. We also eliminated alkali basalts (normative nepheline or kalsilite), especially in data compilations for the Sunda and Honshu arcs, and placed boninitic lavas (.54 wt.% SiO2, ,0.5 wt.% TiO2, plus samples described in the original data sources as boninites) in a group separate from other lavas. A third issue in using large numbers of compiled data is quality. Analytical methods have varied over time, and some labs are more reliable than others. Outliers appear on many of our plots, particularly those involving trace elements. In some cases, where outliers are orders of magnitude from the bulk of the data, we have adjusted axis limits in plots so that outliers are no longer visible but the variation in the bulk of the data is easily seen. However, other than this, we have not made any attempt in this chapter to discriminate between “good” and “bad” data. This approach is deliberately different from other recent reviews of arc data (e.g., Elliott, 2003; Plank, 2003). We are not critical of these other reviews, but we think an alternative, more inclusive approach may be useful until a truly large number of high-quality ICP-MS data become available for a fully representative set of arc magma compositions worldwide. In particular, one focus of this chapter, on the origin of primitive andesites and calc-alkaline magma series, would be all but impossible if we restricted attention to data sets including ICP-MS analyses. We urge readers to be cautious in interpreting our data plots, and to check key points for themselves. Also, we believe this chapter indicates several areas in which additional data would be very valuable. Finally, a substantial limitation of our compilation is that it includes very sparse data on chlorine, fluorine, boron, beryllium, and lithium, and essentially no data on volatile contents (H2O, CO2, sulfur, noble gases). While data on H2O in glass inclusions are beginning to become available, it is not yet clear, for example, to what extent H2O contents in primitive arc magmas

correlate with other compositional characteristics. We look forward to learning more about these topics.

3.18.3 CHARACTERISTICS OF ARC MAGMAS 3.18.3.1 Comparison with MORBs In this section, we compare arc lava compositions (on an anhydrous basis) with compositions of MORB glasses; also see Chapter 3.13. The contrasts are remarkably distinct (Figure 1). Before we go on to describe these contrasts, it is convenient to define “primitive andesites” (lavas with SiO2 . 54 wt.% and Mg# . 60, exclusive of boninites) and “high Mg# andesites” (SiO2 . 54 wt.% and Mg# . 50). These classifications include primitive and high Mg# lavas that are dacites and even rhyolites, as well as true andesites. However, we group them all for brevity. Implicit in our definition of lavas with Mg# . 60 as “primitive,” and lavas with lower Mg# as “evolved” is the assumption that crystallization processes always produce a lower Mg# in derivative liquids, as compared to parental liquids. As far as we know, this assumption is justified on the basis of all available experimental data on crystallization/melting of igneous rocks with Mg# higher than 40 (see compilation in Kelemen (1995)) at oxygen fugacities within 2 log units of Ni-NiO (typical for arcs, Blatter and Carmichael (1998), Brandon and Draper (1996), Gill (1981), and Parkinson and Arculus (1999)), or lower. Oxygen fugacity more than two log units above Ni-NiO may facilitate early and abundant crystallization of FeTi oxides, and thus nearly constant Mg# with decreasing temperature and liquid mass (Kawamoto, 1996). However, this is unlikely beneath most arcs.

3.18.3.1.1

Major elements

We first examine major elements as a function of Mg# (Figure 1). A very small fraction of MORB glasses have Mg# less than 35, whereas lavas with Mg# , 35 are common in arcs. While SiO2 in primitive (Mg# . 60) MORB is restricted to 48 –52 wt.%, primitive arc lavas range from 45 wt.% to more than 60 wt.% SiO2. The contrast for evolved (Mg# , 60) compositions is even more striking. Overall, Mg# versus SiO2 for MORB glasses closely approximates a single liquid line of descent, involving olivine þ plagioclase þ clinopyroxene, with cumulate SiO2 , liquid SiO2 (see Chapter 3.13). The arc lavas show a much broader trend of Mg# versus SiO2, consistent with crystallization of SiO2-poor assemblages (less plagioclase, added hornblende

Characteristics of Arc Magmas and/or FeTi oxides) from a range of parental melts with SiO2 contents of 45 wt.% to more than 60 wt%. SiO2 contents from 45 to greater than 60 wt% are found in arc lavas with a Mg# of 70 (close to Fe/Mg exchange equilibrium with residual mantle peridotite having an olivine Mg# of 90– 91). This is not the result of crystal fractionation from basalts with Mg# , 70. However, with this said, the manner in which these parental melts acquire their differing SiO2 contents is uncertain and controversial (see Sections 3.18.3.2.4 and 3.18.3.2.5).

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TiO2 contents of arc magmas are generally lower than in MORB glasses. TiO2 versus Mg# in the entire arc lava compilation, and in most individual suites, shows a sharp transition from increasing TiO2 with decreasing Mg#, for Mg# . ,50, to decreasing TiO2 with decreasing Mg#, for Mg# , ,50, which is due to fractionation of FeTi oxides from evolved melts. It is probably safe to conclude that primitive arc magmas have not undergone FeTi oxide fractionation, and thus their low TiO2, compared to MORB, is a primary feature or the result of

Figure 1 Molar Mg# versus concentrations of major-element oxides, in weight percent, in arc lavas (left) and MORB glasses (right), both on an anhydrous basis. Outlined fields on the right-hand diagrams show the range of variation in arc lava compilation of Elliott (2003). Arc and MORB data sources described in text. For MORB, squares are EPR and circles are Atlantic, JDF, and Indian Ocean Ridges. For primitive arc lavas with Mg# . 60, circles are basalts and squares are andesites, open symbols are for oceanic arcs, gray symbols are for continental arcs, and black symbols are for the oceanic Aleutian arc. Boninites are shown with small black squares. For evolved lavas with Mg# , 60, open circles with barbs are for oceanic arcs, gray circles with barbs are continental arcs, and small open circles are oceanic Aleutian samples. Inverted gray triangles are for primitive andesites from Mt. Shasta, southern Cascades (Grove et al., 2001). Intra-oceanic arcs in our data set are from Tonga, Kermadec, Bismark/New Britain, New Hebrides, Marianas, Izu-Bonin, South Sandwich, and the Lesser Antilles. Samples from arcs in our compilation which are, or may be, emplaced within older continental material or thick sequences of continentally derived sediment are from the Philippines, Indonesia, Papua New Guinea, SW Japan, NW Japan, Kuriles, Kamchatka, Cascades, Central America, the Andes, and the Greater Antilles. Large filled diamonds are estimated compositions of the continental crust from Christensen and Mooney (1995), McLennan and Taylor (1985), Rudnick and Fountain (1995), and Weaver and Tarney (1984), including Archean estimate of Taylor and McLennan.

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Figure 1 (continued).

magma mixing. Primitive andesites and boninites have the lowest TiO2 contents in our compilation. Al2O3 contents of primitive arc lavas range from 10 wt.% to 19 wt.%, extending to much lower and higher values than primitive MORB glasses. Al2O3 contents of arc lavas increase with decreasing

Mg#, for lavas with Mg# . ,50, for the entire data set and for most individual arc suites. For Mg# less than ,55, Al2O3 decreases with decreasing Mg#, reflecting plagioclase fractionation. This suggests that plagioclase fractionation may play a minor role in differentiation of most arc melts

Characteristics of Arc Magmas

599

Figure 1 (continued).

from Mg# of 70 or more to Mg# of ,55. Instead, olivine, pyroxene, and/or hornblende fractionation may predominate. Given that primitive arc lavas are generally H2O-rich compared to anhydrous

mantle melts at arc Moho depths (e.g., Anderson, 1974; Falloon and Danyushevsky, 2000; Kamenetsky et al., 1997; Macdonald et al., 2000; Pichavant et al., 2002; Roggensack et al., 1997;

600

Geochemistry of Subduction-related Magmatic Arcs

Sisson and Layne, 1993; Sobolev and Chaussidon, 1996), many primary mantle melts are probably in a reaction relationship with olivine (e.g., Mu¨ntener et al., 2001), forming by reactions such as orthopyroxene þ clinopyroxene þ spinel ¼ olivine þ melt. Such melts are in equilibrium with olivine, but will not crystallize olivine upon isobaric cooling, and instead will crystallize websterites (two pyroxene pyroxenites). Other primitive, hydrous arc magmas are olivine saturated; these commonly produce olivine clinopyroxenite cumulates (e.g., Conrad and Kay, 1984; Conrad et al., 1983). In addition, primitive arc magmas probably have temperatures .1,100 8C (e.g., Elkins Tanton et al., 2001; Gill, 1981; Kelemen et al., 2003b), above the thermal stability of hornblende, although this might be uncertain for very H 2O-rich compositions. Thus, pyroxenite (clinopyroxenite, websterite) fractionation is most likely responsible for the increase in Al2O3 with decreasing Mg#, for liquid Mg# of more than 70 to ,55 (e.g., Conrad and Kay, 1984; Conrad et al., 1983). Pyroxenites have seismic velocities similar to or only slightly less than residual mantle peridotites. The temperature and melt content of the sub-arc mantle are poorly constrained, and seismologists commonly report sub-Moho P-wave velocities less than 8 km s21 (e.g., Fliedner and Klemperer, 1999; Holbrook et al., 1999; Suyehiro et al., 1996). Thus, the igneous crust may extend well below the seismic Moho in arcs. However, by the same token, seismic data certainly do not require abundant pyroxenite. Pyroxenites, while well represented among arc plutonic xenoliths, comprise a very small proportion of exposed arc crustal sections (e.g., DeBari and Coleman, 1989; Miller and Christensen, 1994). Thus, pyroxenites may be removed by viscous “delamination” during or after arc magmatism (e.g., Arndt and Goldstein, 1989; DeBari and Sleep, 1991, 1996; Herzberg et al., 1983; R. W. Kay and S. M. Kay, 1988, 1991, 1993; S. M. Kay and R. W. Kay, 1985; Turcotte, 1989). We will return to this topic in Sections 3.18.4.2 and 3.18.5.1. Rare primitive arc lavas, including well-studied compositions that almost certainly represent liquid compositions, have Al2O3 higher than primitive MORB. These are generally interpreted as products of equilibration of nearly anhydrous basaltic melts with residual mantle peridotite at 1– 1.2 GPa, just beneath the base of arc crust in the Cascades and Indonesia (e.g., Bartels et al., 1991; Elkins Tanton et al., 2001; Sisson and Bronto, 1998), although see also Turner and Foden (2001). Although these compositions are not common (Figure 1), they are nonetheless of great importance because well-known phase equilibria for nearly anhydrous melt/mantle equilibration allows these lavas to be used to

place fairly tight constraints on sub-arc temperature at specific depths in the mantle wedge (e.g., Elkins Tanton et al., 2001). Al2O3 contents in evolved arc magmas range up to more than 20 wt.%, much higher than in evolved MORB. Interpretation of Al2O3 contents of evolved arc lavas is notoriously difficult, because of the potential for incorporation of accumulated plagioclase in porphyritic lavas (Brophy, 1989; Crawford et al., 1987). However, some carefully studied “high alumina basalts” are probably liquid compositions (Baker and Eggler, 1983, 1987; Sisson and Grove, 1993a,b). In addition, some authors have maintained that high Al2O3 arc lavas with Mg# , 0.5 might represent primary melts from diapirs of subducted basalt that rise into the mantle wedge, but do not chemically equilibrate with residual mantle peridotite (e.g., Brophy and Marsh, 1986; Johnston and Wyllie, 1988; Marsh, 1976; Myers et al., 1986a,b, 1985). This topic is visited again in Section 3.18.3.2.3. Alternatively, moderate Mg# lavas may be derived via crystal fractionation from primitive melts, or may be partial melts of arc lower crust. From this perspective, Al2O3 increases due to crystallization of plagioclase-free cumulates, where plagioclase saturation is suppressed by abundant H2O (greater than ,2 wt.% H2O) in primitive arc melts (e.g., Baker and Eggler, 1983, 1987; Kelemen et al., 1990a; Mu¨ntener et al., 2001, Sisson and Grove, 1993a,b). In this review, using compiled data on arc lava compositions, it is impossible to improve upon previous estimates of oxygen fugacity or Fe3þ/Fe2þ in arc magmas (generally close to NiNiO, e.g., Blatter and Carmichael (1998), Brandon and Draper (1996), Gill (1981), and Parkinson and Arculus (1999)). Thus, here we concentrate on total Fe as FeO (FeOt). Both primitive and evolved arc magmas extend to FeOt contents much lower than in MORB glasses. While some of this difference among evolved compositions results from FeTi oxide fractionation in relatively oxidizing arc magmas, the differences between primitive arc lavas and primitive MORB are clear, and are not attributable to FeTi oxide fractionation because FeTi oxides are not saturated in primitive arc lavas. Instead, this reflects a difference in the FeOt content of primary magmas, and/or the effects of magma mixing. MgO contents among arc lavas show a much broader range than in MORB glasses. Arc lava MgO contents higher than MORB glasses may be due, in part, to incorporation of compositions with accumulated clinopyroxene and/or olivine into our data set. However, boninite suites clearly include high MgO liquids (e.g., Falloon and Green, 1986; Sobolev and Danyushevsky, 1994), and some picritic and ankaramitic arc lava

Characteristics of Arc Magmas

3.18.3.1.2

We are cautious about fractionation correction of major elements

Na2O concentration in basalts, corrected to 6 wt.% MgO, has been used as an indicator of the relative degree of mantle melting in the arc mantle source (e.g., Plank and Langmuir, 1988, 1993). This method involves two explicit assumptions. First, it is assumed that that primary melts have a common MgO content. While it is beyond the scope of this chapter to investigate this assumption on a volcano-by-volcano basis, in Figure 2, we show that this assumption seems questionable for Aleutian lavas, which show a wide variation in MgO at high Mg#. Lavas from Piip volcano with 6 wt.% MgO are primitive, so that their Na2O contents may be close to those of the “primary” magma, whereas lavas from Okmok and Seguam volcanoes with 6 wt.% MgO are fractionated, with Na2O contents higher than in corresponding, primary melts in

7 6

Na2O (wt.%)

5 4 3 2 1 0 14 12

all data Okmok

Seguam

10 MgO (wt.%)

compositions are also believed to be liquid compositions (e.g., Eggins, 1993; Nye and Reid, 1986; Ramsay et al., 1984). Primitive arc lavas also extend to MgO contents much lower than primitive MORB glasses. These have MgO too low to be derived from high MgO, primitive arc lavas via olivine fractionation. Alternatively, primitive andesites could be derived via olivine fractionation from picritic or komatiitic primary melts. However, as we will show in Sections 3.18.3.2 and 3.18.3.3, primitive andesites generally have trace-element characteristics that are distinct from primitive basalts, so the two cannot be related by crystal fractionation alone. Instead, as for low FeOt, low MgO in primitive arc andesites reflects either low MgO in primary melts equilibrated with residual mantle peridotite, or the effects of magma mixing. Like FeO and MgO, both primitive and evolved arc lavas extend to CaO contents much lower than in primitive and evolved MORB glasses, and andesites are the low CaO end-member among primitive lavas. It is ironic that low CaO is one of the characteristics of these, end-member “calcalkaline” lavas; this suggests that the “tholeiitic” versus “calc-alkaline” terminology should be changed. Na2O, K2O and, to a lesser extent, P2O5 in both primitive and evolved arc lavas extend to higher and lower concentrations than in primitive and evolved MORB glasses. Again, primitive andesites stand out, having the highest Na2O, K2O, and P2O5 among primitive compositions. And, as for the other elements discussed above, these characteristics are not due to crystal fractionation, but must be characteristics of primary magmas or must arise via magma mixing.

601

Piip

8 6 4 2 0

0

10

20

30 40 50 Molar Mg#

60

70

80

Figure 2 Molar Mg# versus wt.% Na2O for Aleutian lavas (compiled by Kelemen et al., 2003b), illustrating that variation trends have different slopes for different volcanoes, and that these trends cross. As a result, it would be unwise to infer Na2O concentrations in primary Aleutian magmas using an “average” fractionation trend and the compositions of evolved lavas. Data for Seguam volcano mainly from Singer et al. (1992a, b) and data compiled by James Myers and Travis McElfrish and available at http://www.gg.uwyo.edu/ aleutians/index.htm. Data for Okmok volcano mainly from Class et al. (2000), Miller et al. (1992, 1994), Nye (1983), and Nye and Reid (1986, 1987). Okmok data with Mg# , 80 omitted. Data for Piip volcano from Yogodzinski et al. (1994). Light grey circles, all Aleutian data (compiled by Kelemen et al., 2003). Other shapes are for individual volcanoes, as shown in legend. Darker symbols used in fits (linear in top panel, exponential in bottom panel); open symbols omitted from fits.

equilibrium with mantle olivine. Because the trends of Na2O versus Mg# for Okmok and Seguam volcanoes cross, correction to 6 wt.% MgO yields higher Na2O in Okmok compared to

602

Geochemistry of Subduction-related Magmatic Arcs

Seguam, whereas primary Okmok melts (Mg# , 70) probably have lower Na2O compared to primary Seguam melts. Second, it is assumed that Na2O in arc magmas is derived primarily from the mantle, without a significant contribution from subducting sediment and oceanic crust. We think this also is uncertain, because fluids and melts in equilibrium with subducting, eclogite facies metasediment and metabasalt contain abundant dissolved sodium. This dissolved sodium could be carried into the mantle wedge together with H2O and other components derived from subducting material (see Chapter 3.17, and references therein). For example, primitive andesites in the Aleutians that have exceptionally high Na2O also have exceptionally high La/Yb, Sr/Nd, Dy/Yb, and 143 Nd/144Nd, and exceptionally low lead and strontium isotope ratios (Kelemen et al., 2003b), which likely reflects incorporation of an eclogite melt component. We will return to this point in Section 3.18.3.2. If partial melts of eclogite have ,5 wt.% Na2O (Rapp et al., 1999; Rapp and Watson, 1995), addition of 5% eclogite melt to some part of the mantle wedge adds 0.25 wt.% Na2O to that region. This can be compared to, e.g., 0.4 wt.% Na2O in fertile “pyrolite” (Ringwood, 1966) or #0.05 wt.% in depleted oceanic peridotites (Dick, 1989). Thus, it is plausible that a “slab melt” component might add significant and potentially variable amounts of sodium to the mantle “source” of arc magmas, particularly if the arc mantle has been previously depleted by melt extraction beneath a mid-ocean ridge and/or a back-arc basin. In order to more quantitatively constrain the relative contributions of sodium from melt or fluid derived from subducting material, versus the pre-existing peridotite in the mantle wedge, it would be useful to know the degree of depletion of the pre-existing peridotite, due to previous melt extraction events prior to arc magmatism. One can try to use the concentration of “immobile” elements such as niobium, together with mantle/melt distribution coefficients, to estimate the degree of depletion of the mantle “source” of each arc magma, assuming that “immobile” elements are not added by fluids or melts of subducting material (Langmuir, personal communication, 2003). However, in addition to sodium, eclogite melts also contain ,5 ppm niobium (Rapp and Watson, 1995), while primitive mantle is estimated to contain ,1 ppm niobium (Hofmann, 1988; Sun and McDonough, 1989). Therefore, determining the extent of prior depletion in the mantle wedge is also problematic for systems open to eclogite melt. Third, use of Na2O as an indicator of mantle melting processes involves the implicit

assumption that Na2O contents in arc magmas have not been affected by open-system processes in the crust, such as magma mixing and assimilation. Again, this assumption is violated in some cases, particularly when considering evolved, rather than primitive, lava compositions (e.g., Grove et al., 1982, 1988; Hildreth and Moorbath, 1988; McBirney et al., 1987).

3.18.3.1.3

Distinctive, primitive andesites

As will be seen throughout this chapter, despite limited data, primitive andesites define an end-member on almost all compositional variation diagrams. The primitive andesite end-member is distinct from primitive MORB glasses in ways that epitomize the overall difference between arc lavas and MORB (Table 1 and Figure 3). High Mg# andesites, and their plutonic equivalents, are end-member calc-alkaline lavas, as distinct from the tholeiitic magma series. Here we define calc-alkaline magmas as having both high Na þ K at high Mg# (Irvine and Baragar, 1971) and high SiO2 at high Mg# (i.e., high SiO2 at low Fe/Mg, Miyashiro, 1974); tholeiitic magmas have lower SiO2 and Na þ K at the same Mg#, when compared to calc-alkaline magmas. (Please note that in this chapter we do not consider alkaline lavas, i.e., nepheline or kalsilite normative compositions) While primitive basalts, and evolved, tholeiitic lavas, are found in a variety of plate tectonic settings, it is plain that calc-alkaline andesite lavas are found almost exclusively in arcs (Gill, 1981). Thus, one could argue that the genesis of primitive andesite is the defining process of arc magmatism. While high Mg# andesites are clearly less voluminous than tholeiitic lavas in most arcs (Figure 4; see also, e.g., White and McBirney (1978)), we consider them very important in other ways. Explaining the difference between calcalkaline and tholeiitic magma series has been one of the central topics of igneous petrology for almost a century (e.g., Baker et al., 1994; Bowen, 1928; Brophy, 1987; Fenner, 1929, 1937; Green, 1976; Grove et al., 1982, Green and Ringwood, 1967, 1966; Grove and Kinzler, 1986; 2001; Kay, 1978, 1980; S. M. Kay and R. W. Kay, 1994; Kay et al., 1982; Kelemen, 1986, 1990; Kuno, 1950, 1968; Kushiro, 1969, 1974; Kushiro and Yoder, 1972; McBirney et al., 1987; Miller et al., 1992; Nicholls and Ringwood, 1973; Nicholls, 1974; Osborn, 1959; Sisson and Grove, 1993a,b; Tatsumi, 1981, 1982; Tatsumi and Ishizaka, 1981, 1982; Wilcox, 1944), and there is still no community-wide consensus on this. Similarly, the estimated bulk composition of the continental crust (e.g., Christensen and Mooney, 1995; McLennan and Taylor, 1985;

Table 1(a)

Average primitive MORB and arc basalts (Molar Mg# . 60).

MORB

Average oceanic

Average continental

Kermadec

Lesser Antilles

Marianas

New Hebrides

Scotia

Tonga

Aleutian

Andean

Cascades

Central America

Greater Antilles

Honshu

Kamchatka

Luzon

n SiO2 TiO2 Al2O3 FeO(T) MnO MgO CaO Na2O K2O P2O5

203 50.51 1.22 15.97 8.85 0.16 8.57 11.85 2.57 0.16 0.14

503 50.46 0.91 15.72 8.52 0.17 9.84 11.44 2.35 0.45 0.15

497 51.33 0.98 15.70 8.72 0.17 9.48 9.93 2.61 0.88 0.22

36 51.12 0.81 15.65 8.77 0.16 9.07 11.50 2.27 0.53 0.12

84 48.25 0.85 14.48 9.17 0.17 13.69 10.79 2.01 0.45 0.15

168 51.04 1.01 16.64 8.02 0.15 8.57 11.51 2.62 0.31 0.13

65 50.26 0.69 13.67 8.74 0.19 11.33 11.96 2.06 0.90 0.18

41 51.50 1.06 16.97 8.05 0.15 7.74 11.34 2.72 0.32 0.15

70 50.57 0.94 15.59 8.60 0.17 9.39 11.99 2.16 0.43 0.16

66 50.50 0.79 16.51 8.58 0.16 9.22 11.09 2.39 0.61 0.14

56 52.58 1.03 16.66 8.11 0.15 8.28 8.80 2.97 1.19 0.21

60 51.62 0.87 16.83 7.89 0.15 9.12 9.86 2.93 0.54 0.18

78 50.27 1.02 14.60 9.50 0.17 10.63 10.49 2.31 0.77 0.23

21 50.23 1.04 15.54 9.20 0.20 10.57 9.57 2.50 0.91 0.23

137 51.13 1.03 16.14 8.79 0.16 9.29 9.68 2.53 0.99 0.26

78 52.22 0.92 14.76 8.78 0.17 9.39 9.96 2.73 0.87 0.19

24 50.85 0.78 14.20 8.92 0.18 10.86 10.75 2.24 0.99 0.22

Molar Mg#

63.23

65.24

64.48

70.97

65.05

69.53

64.12

67.17

64.49

65.00

65.09

11.85 25.87 2.85 14.88 3.43 1.07 3.55 0.51 3.32 0.68 1.95 0.29 1.82 0.28 32.51 246.59 397.96 41.20 158.74 91.91 81.30 18.63 425.70 18.69 92.70 6.23 0.71 295.04 2.14 0.45 3.36 2.03 0.53 0.70401 0.51292 18.75 15.57 38.46

n n 159 6.69 10 157 15.27 10 65 152 10.76 10 155 2.61 10 157 0.88 10 107 122 0.48 10 91 65 84 63 159 1.88 10 138 0.30 10 110 35.80 10 107 255.10 10 145 317.10 10 80 146 110.00 10 40 82.40 10 68 77.60 10 140 8.30 10 153 274.90 10 141 20.40 10 144 65.90 10 135 2.30 10 87 155 124.00 10 133 1.73 10 83 0.26 10 83 2.80 10 139 1.11 10 102 0.28 8 133 0.70419 19 104 0.51293 11 76 18.84 2 76 15.62 2 75 38.75 2

8.72 18.94 0.72 10.93 2.68 0.91 2.98 0.43 2.93 0.40 1.69 0.20 1.59 0.25 34.76 235.38 974.55 61.22 442.69 78.21 74.59 10.22 314.97 17.45 59.65 4.79 0.32 137.67 1.54 0.26 2.40 2.96 1.07 0.70482 0.51291 19.30 15.71 38.90

n n 50 5.00 46 56 11.73 51 7 55 9.10 40 57 2.72 40 56 1.01 51 45 3.47 36 18 0.55 16 41 3.86 37 10 43 2.38 37 11 56 2.29 40 42 0.37 7 44 33.64 15 39 245.57 21 47 322.50 22 13 34.79 14 52 128.30 20 34 40 60.08 12 57 5.61 50 57 231.78 51 54 24.14 21 53 65.14 22 52 3.81 14 12 0.08 10 54 65.47 48 13 1.61 14 11 0.14 12 32 1.08 26 38 0.41 12 39 0.17 11 46 0.70303 45 42 0.51306 41 26 18.47 30 26 15.52 30 26 38.17 30

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Sc V Cr Co Ni Cu Zn Rb Sr Y Zr Nb Cs Ba Hf Ta Pb Th U 87 Sr/86Sr 143 Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb

4.13 11.46 1.84 9.30 2.96 1.13 4.04 0.68 4.70 0.93 3.00 0.47 2.72 0.41 36.75 245.77 357.10 41.76 135.24 69.87 67.27 2.93 141.42 27.48 92.63 6.04 0.05 30.70 2.31 0.31 0.38 0.28 0.23 0.70274 0.51310 18.38 15.50 38.01

66.29 n 59 62 6 60 66 66 27 42 25 6 29 4 62 40 45 39 48 33 46 30 30 33 55 46 54 20 6 44 41 28 14 34 26 104 90 86 86 86

7.01 15.67 2.11 10.14 2.70 0.95 3.10 0.52 3.31 0.62 2.00 0.28 1.86 0.27 36.37 254.01 575.68 44.17 239.67 84.86 72.22 9.89 306.74 19.47 62.21 3.99 0.32 132.96 1.65 0.24 2.48 1.52 0.59 0.70389 0.51298 18.80 15.59 38.49

n 168 181 55 168 172 181 134 97 133 58 136 33 171 97 112 119 132 56 135 74 97 179 181 145 145 129 54 175 87 64 109 111 105 141 124 75 75 75

8.38 19.26 2.51 11.82 2.79 0.90 2.86 0.47 2.77 0.58 1.65 0.24 1.57 0.23 38.00 336.00 420.11 44.67 131.72 126.00 72.17 14.83 499.81 17.86 51.58 1.83 0.14 257.24 1.58 0.16 4.12 1.19 0.41 0.70392 0.51300 18.40 15.52 38.31

63.06 n 21 20 21 21 21 21 20 20 21 20 21 6 21 6 14 16 18 6 18 6 6 20 21 21 21 20 6 21 18 6 10 20 20 4 4 4 4 4

6.49 15.19 2.28 9.88 2.86 1.03 3.34 0.61 3.87 0.83 2.33 0.36 2.27 0.35 34.45 237.00 266.33 35.94 84.67 65.30 77.06 7.24 202.00 23.29 86.22 6.71 0.11 84.23 2.15 0.50 1.63 0.89 0.23 0.70337 0.51301 18.99 15.62 38.78

65.81 n 14 16 11 16 16 16 15 11 16 11 16 11 16 11 11 11 15 10 15 10 16 16 16 16 16 16 11 16 11 11 12 12 11 14 5 2 2 2

64.96

8.95 18.11

n 12 12

10.83 2.74 1.00 3.49 0.49 3.41 0.58 2.20

12 12 11 5 9 6 5 6

1.62 12 0.19 8 44.58 9 222.60 10 716.67 9 45.07 6 182.22 9

24.96 451.64 16.64 50.91 5.88 1.34 274.91 1.53 0.04 5.99 0.78 0.51 0.70406 0.51292 18.89 15.55 38.58

10 11 11 11 6 5 11 9 5 8 8 5 7 3 2 2 2

n 27 27 5 27 27 27 12 20 25 5 11

n n 5.98 18.79 28 11.29 24 14.39 41.23 26 24.22 24 2.21 7.10 10 1.90 5 8.80 20.60 26 13.74 24 2.33 4.37 28 3.16 24 0.80 1.24 28 1.02 24 2.28 4.79 10 3.37 11 0.43 0.60 26 0.49 20 4.04 10 2.93 5 2.81* 0.52 0.75 10 0.65 5 1.36 2.14 10 1.79 5 0.31 9 0.29 5 1.40 27 1.83 28 1.88 24 0.21 19 0.28 26 0.27 24 38.71 20 26.54 28 34.24 5 294.38 8 198.36 22 224.00 5 449.93 27 344.42 26 358.75 24 44.09 16 41.06 9 43.08 22 130.08 26 130.49 28 151.29 24 78.50 4 46.80 5 91.86 7 83.86 22 63.00 5 9.88 21 29.23 28 16.42 23 445.09 23 532.34 28 469.38 24 12.90 18 19.74 28 20.91 22 54.94 18 118.96 28 105.96 24 4.05 17 11.76 24 6.70 21 0.46 21 1.00 19 0.38 22 195.54 26 317.81 28 271.33 24 1.58 22 2.76 26 2.22 22 0.25 12 0.99 20 0.48 17 3.95 10 7.21 19 2.26 10 1.21 23 3.77 26 1.33 22 0.58 19 1.35 11 0.39 22 0.70315 19 0.70515 11 0.70382 27 0.51303 16 0.51262 6 0.51285 27 18.69 14 18.50 5 18.94 26 15.53 14 15.61 5 15.62 26 38.20 14 38.73 5 38.61 26

67.64

14.75 30.54 1.35 15.83 3.42 1.08 3.46 0.43 3.29 0.60 1.88 0.25 1.68 0.25 33.20 268.53 491.51

n 29 29 17 29 29 29 29 17 29 17 29 17 29 17 13 29 29

245.05 118.17

29 13

10.63 437.96 18.67 79.94 7.03 0.22 315.53 1.53 0.29 1.40 1.08 0.37 0.70388 0.51301 19.02 15.56 38.68

29 29 29 29 29 10 29 25 17 26 26 26 25 25 21 21 21

5.60 14.06 1.70 9.88 2.80 0.93 3.09 0.51 2.39 0.74 2.15 0.32 1.95 0.29 33.21 253.20 269.13 141.53 108.86 81.89 12.48 284.60 13.15 63.38 3.51 0.43 284.88 1.70 3.20 0.75 0.70432 0.51293 19.99 15.73 39.27

n n 16 16 6 16 16 16 6 16 6 6 6 6 16 16 14 15 186.65 20 15 507.53 45 57.13 15 15 191.52 57 7 9 12 22.10 39 16 715.20 38 16 22.54 21 16 122.55 22 14 11.80 21 2 16 498.13 24 15 5 15 1 0.70437 27 1 0.51277 9 1 1 1

6.92 16.89 2.64 12.60 3.33 1.09 3.49 0.51 3.51 0.70 2.02 0.29 1.89 0.30 35.33 260.45 442.26 41.72 135.08 75.00 83.52 14.45 345.68 19.54 87.14 2.90 0.49 279.70 2.39 0.13 2.77 0.88 0.43 0.70344 0.51307 18.26 15.49 37.94

66.05 n 41 41 23 41 41 39 41 27 31 23 24 22 41 40 40 31 39 36 40 9 27 33 41 32 32 32 27 40 32 22 22 32 31 28 26 21 21 21

15.99 32.76

n 13 13

18.26 2.80 0.94 3.38 0.38 2.76

12 9 13 8 9 8

1.62

8

1.38 0.22 36.32

13 7 5

339.57 35.63 96.57

7 7 7

37.07 566.18 16.10 96.00 5.05

11 11 10 11 11

335.80 12 1.88 6 0.16 5 5.80 0.62 0.70442 0.51263 18.20 15.51 38.14

13 7 4 5 1 1 1

Major element oxides in wt.%, trace elements in ppm. All analyses are included in major element average. Analyses with values for at least four of the REEs are included in the trace element averages, and only if there are ten of these analyses, is an individual arc included in the table. Only elements for which at least five analyses are available are included in the table.p For Aleutian samples without Dy analyses, we estimated Dy , 7.12pTb

604

Geochemistry of Subduction-related Magmatic Arcs Average primitive arc andesites (molar Mg# . 60).

Table 1(b)

n majors SiO2 TiO2 Al2O3 FeO(T) MnO MgO CaO Na2O K2O P2O5 Molar Mg#

Continental

Oceanic

142 58.05 0.79 15.96 6.14 0.12 6.56 7.20 3.31 1.67 0.22 65.18

32 57.72 0.64 15.16 6.69 0.14 7.95 7.32 2.95 1.27 0.14 66.29 n

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu Li Be Sc V Cr Co Ni Cu Zn Rb Sr Y Zr Nb Cs Ba Hf Ta Pb Th U 87 Sr/86Sr 143 Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb

18.89 37.44 5.38 20.89 3.92 1.08 3.92 0.51 3.09 0.55 1.63 1.54 0.23 6.73 0.87 20.98 158.27 326.83 31.36 137.89 91.75 74.43 45.66 586.66 17.13 137.19 7.94 2.27 501.74 3.56 0.85 8.45 4.51 1.57 0.70469 0.51277 18.53 15.56 38.36

59 53 13 42 56 59 26 47 23 15 22 57 51 3 3 41 19 55 36 52 8 14 47 48 37 41 34 20 41 32 22 20 35 29 31 27 11 11 11

Aleutian 47 59.03 0.69 16.61 5.22 0.10 5.65 7.35 3.64 1.50 0.20 65.62 n

26.14 196.66 260.75 30.69 118.63 75.44 69.02 29.52 358.80 20.34 91.04 4.08 1.61 273.27 1.83 0.08 6.00 3.90 0.37 0.70493 0.51288 18.81 15.58 38.54

Boninites

18 21 21 8 22 18 18 25 25 25 25 20 4 27 5 4 13 16 6 14 6 4 4 4

348 56.83 0.25 13.22 7.93 0.15 10.64 8.35 2.02 0.56 0.06 69.50 n

16.02 37.27

28 28

20.20 3.95 1.12 3.61 0.43 3.06*

28 28 28 5 23 28

1.77 1.32 0.19

7 28 27

17.55 170.00 252.76 22.07 95.14 63.67 72.33 20.52 1,035.88 14.85 115.33 4.95 0.44 309.62 3.06 0.25 4.70 1.99 0.82 0.70291 0.51308 18.30 15.47 37.76

25 2 27 21 28 3 3 21 27 22 15 21 23 28 23 20 13 23 24 13 18 18 18 18

n 1.88 4.44 0.92 2.79 0.77 0.28 0.92 0.17 1.01 0.31 0.70 0.80 0.14 7.61 0.26 36.45 188.01 696.03 43.10 191.76 65.28 57.13 9.47 141.84 7.59 39.05 2.17 0.19 54.53 0.70 0.13 1.83 0.42 0.26 0.70423 0.51294 18.68 15.53 38.31

74 73 23 63 71 76 55 38 60 21 55 69 53 14 16 54 70 76 38 75 25 31 76 77 68 74 31 14 73 43 22 15 28 19 55 50 37 37 37

Major element oxides in wt.%, trace elements in ppm. Trace-element averages are calculated as in Table 1(a), except for oceanic andesites, where REE averages are not calculated (too few and variable analyses) and where all analyses are included for other trace elements.

Rudnick and Fountain, 1995; Weaver and Tarney, 1984) is almost identical to some high Mg# andesites in both major and trace-element concentrations, and some authors have proposed that the genesis of continental crust involved

processes similar to the generation of high Mg# andesites today (e.g., Defant and Kepezhinskas, 2001; Drummond and Defant, 1990; Ellam and Hawkesworth, 1988b; Kelemen, 1995; Kelemen et al., 1993, 2003b; Martin, 1986, 1999; Rapp and

Characteristics of Arc Magmas MORB glasses Kermadec Tonga Marianas New Hebrides Scotia Lesser Antilles Greater Antilles Aleutian Kamchatka Andean Cascades Central America

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Figure 3 Extended trace-element diagrams for average arc lavas (Table 1). Concentrations are normalized to NMORB (Hofmann, 1988). Primitive arc basalts are remarkably similar from one arc to another, and consistently distinct from MORB. In the oceanic Aleutian arc, and in continental arcs, primitive andesites are more enriched than primitive basalts. For plotting purposes some REE abundances are extrapolated from neighboring REEs with more analyses (Pr in Lesser Antilles, Dy in Greater Antilles, Er in Aleutian).

Watson, 1995; Taylor, 1977). Again, however, there is no consensus on this. Finally, plutonic rocks with high Mg# andesite compositions probably form the bulk of the major calc-alkaline plutons in orogenic belts, such as the Mesozoic batholiths along the Pacific margins of

North and South America. For example, the average composition of the Peninsular Ranges batholith in southern California is essentially identical to that of continental crust (Gromet and Silver, 1987; Silver and Chappell, 1988). Similarly, the average composition of exposed, Eocene

Aleutians, west of Adak

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Characteristics of Arc Magmas to Miocene plutonic rocks in the Aleutian arc is that of high Mg# andesite (Kelemen et al., 2003b). Thus, high Mg# andesite magmas may be more commonly emplaced as plutonic rocks in the middle and upper crust (Kay et al., 1990; Kelemen, 1995), and may be under-represented among erupted lavas. For these reasons, and because this topic has not received a recent, comprehensive review, one emphasis in this chapter is documentation of the difference between primitive arc basalts and andesites, and evidence bearing on its origin.

3.18.3.1.4

Major elements in calc-alkaline batholiths

…there is no reason to suppose that the relative amounts of magmas of different compositions erupted on the surface should be proportional to their amounts…at depth. Kuno (1968, p. 168)

This section is based on a limited compilation of plutonic rock compositions from “intermediate, calc-alkaline batholiths,” such as are common among Mesozoic and early Tertiary exposures associated with circum-Pacific arcs. Section 3.18.4 describes the composition of plutonic rocks from exposed arc sections that extend into the lower crust, including a much larger proportion of mafic gabbros. Our data include the few data available in GeoRoc, our previous compilation of Aleutian plutonic rock compositions (Kelemen et al., 2003b), plutonic rocks from the Tanzawa complex, interpreted as accreted mid-crustal rocks from the Izu – Bonin arc (Kawate and Arima, 1998), and a compilation of limited data from highly “calc-alkaline” batholiths such as the Mt. Stuart and Chilliwack batholiths in the North Cascades (Erikson, 1977; Kelemen and Ghiorso, 1986; Tepper et al., 1993), the Peninsular Ranges batholith in Baja California (Gromet and Silver, 1987; Larsen, 1948; Silver and Chappell, 1988), and the Ladakh batholith in northwestern India (Honegger et al., 1982). Unfortunately, because this data set is small, it is not clear to what extent the compiled compilations are representative of intermediate arc plutons in general. A second problem, in interpreting plutonic rock compositions, is determining the extent to which

607

they represent liquid versus “cumulate” compositions, where “cumulate” is taken to mean a component formed by partial crystallization of a melt, after which the remaining melt was extracted from the system of interest. A wide array of possible plutonic compositions can be envisioned, lying between these two extremes. For example, some plutonic rocks may be cumulates plus a small amount of “trapped melt.” Others may be cumulates affected by interaction with unrelated, migrating melts. Cumulates with abundant plagioclase should generally have high Sr/Nd, since strontium is much more compatible than neodymium in plagioclase, and Eu/Sm, since europium is generally much more compatible than samarium in plagioclase (depending on oxygen fugacity). In general, plutonic rocks with more than 55 –60 wt.% SiO2 closely resemble liquid compositions in many ways, often containing abundant incompatible elements and lacking anomalously high Sr/Nd and Eu/Sm. We quantify this for intermediate to felsic plutonic rocks from the Jurassic Talkeetna arc section in Section 3.18.4.1.1. Our compiled plutonic rock compositions are generally more SiO2-rich, at a given Mg#, than the compiled arc lavas (Figure 5). This is particularly clear in comparing Aleutian lavas to Aleutian plutons (Kelemen et al., 2003b). TiO2 is highest in plutonic rocks with Mg# of ,60. Al2O3 is low in plagioclase-poor, high Mg# pyroxenites, and then—generally—similar to the lower Al2O3 arc lavas at a given Mg#. Arc plutons have lower FeO, MgO, and CaO, at a given Mg#, than the bulk of arc lavas. Although some primitive cumulates in our compilation have very low alkali contents, in general Na2O contents in arc plutons and lavas are comparable, while K2O contents are generally higher in arc plutons compared to arc lavas at the same Mg#. To summarize, in all of their major-element characteristics, the samples in our compilation of arc plutons are more strongly calc-alkaline, and include more high Mg# andesite compositions, than typical arc lavas. It seems that high Mg# andesite liquids may be better represented among intermediate plutonic rocks than among arc lavas. This is certainly the case for exposed Aleutian plutons compared to Aleutian lavas. In order to explain this, following Kay et al. (1990), we have suggested that this

Figure 4 Histograms of wt.% SiO2, wt.% K2O, and 207Pb/204Pb for intra-oceanic arc lavas in our compilation. SiO2 and K2O are for lavas with Mg# . 50, while Pb isotopes are for all lavas. This diagram shows that relatively low K basalts predominate over relatively K-rich, high Mg# andesites in all intra-oceanic arcs except the western Aleutians. The western Aleutian arc also has the least radiogenic Pb isotopes of any intra-oceanic arc. Thus, the predominance of primitive andesites in the western Aleutians is probably not due to recycling of components from subducting, continental sediment, nor to crustal contamination involving pre-existing continental material.

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Geochemistry of Subduction-related Magmatic Arcs difference arises due to the relatively high viscosity of intermediate to felsic magmas as they lose H2O by degassing in the mid- and upper crust (Kelemen, 1995; Kelemen et al., 2003b). Lower H2O, low-SiO2 basaltic melts continue to rise to the surface and erupt, whereas higher SiO2 magmas, with initially higher H2O, rise more slowly and crystallize faster.

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…primitive magnesian andesites and basaltic andesites from the Mt. Shasta region, N. California…form by hydrous mantle melting Grove et al. (2003)

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[A few scientists] seem to want to cling to the possibility that andesites are primary mantle-derived melts despite overwhelming evidence from experimental petrology, trace-element geochemistry, mineral chemistry, petrography, textural, and field relations to the contrary. It would be hard to find very many students of arc petrology who would argue…that andesites, even Mg-rich ones, are… primary mantle melts. Anonymous (2003, personal communication)

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Figure 5 Histograms of wt.% SiO2 for arc lavas and a limited compilation of samples from arc plutons and batholiths. In general, arc plutonic rocks have higher SiO2 content at similar Mg#. Data sources as in Figure 1.

We turn now to the chemical characteristics of primitive arc lavas (Mg# . 60). First, we examine major-element variation as a function of Na2O and TiO2 contents. Although Na2O contents of primitive basalts and andesites overlap, plots of Na2O versus TiO2, FeO, MgO, and CaO clearly discriminate between boninites (very low TiO2 and Na2O) primitive basalts (high TiO2, FeO, MgO, and CaO at a given Na2O), and primitive andesites (low TiO2, FeO, MgO, and CaO at a given Na2O) (Figure 6). It is evident from the trends of these elements versus Mg# (Figure 1) that these variations do not arise from crystal fractionation. For example, arc lavas with Mg# of ,70 or more, in Fe/Mg equilibrium with mantle olivine having Mg# of 90 – 91 or more, have SiO2 contents ranging from 45 wt.% to 63 wt.%. These cannot be related by fractional crystallization. The antithetical behavior of sodium and titanium (high sodium and low titanium in primitive andesites, high titanium and low sodium in primitive basalts; ,2 wt.% Na2O in both primitive MORB and primitive arc basalts, but 0.3 –1 wt.% TiO2 in primitive arc basalts compared to 1.2 wt.% in primitive MORB, Table 1) suggests that sodium contents of arc magmas may not be a good indicator of the degree of partial melting in the sub-arc mantle. (Also see Figure 2, and the last three paragraphs of Section 3.18.3.1.1.) Although there is some overlap, plots of rubidium, barium, thorium, strontium, lead, zirconium, hafnium, and light REE versus Na2O also

Characteristics of Arc Magmas 70

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Figure 6 Wt.% Na2O versus other major-element oxides, in wt.%, plus molar Ca/(Ca+Na) versus ppm Ti/Zr, for primitive arc lavas (Mg# . 60). Many of these plots clearly show distinct compositional fields for primitive basalts, primitive andesites, and boninites. While most of the primitive andesites are from “continental” arcs, they plot together with western Aleutian primitive andesites, which are from an intra-oceanic arc and have MORB-like Sr, Pb, and Nd isotope ratios. Thus, assimilation of older, continental material is not essential to producing the distinctive composition of primitive andesites. Large filled circles show values for average MORB glasses from the East Pacific Rise, Juan de Fuca Ridge, and Indian Ocean. Other symbols and data as for Figure 1.

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Geochemistry of Subduction-related Magmatic Arcs

discriminate between these groups (for example, lanthanum versus Na2O, Figure 7), with high LILEs, light REEs, and other highly incompatible elements in primitive andesites as compared to primitive basalts at a given Na2O concentration. Middle REE concentrations are similar in both andesites and basalts, while heavy REE and related elements (holmium, erbium, thulium, ytterbium, lutetium, yttrium, scandium, vanadium) are lower in andesites than in basalts at a given Na2O concentration (e.g., scandium and ytterbium versus Na2O, Figure 7). Low heavy REE and titanium contents in arc magmas have sometimes been considered to be indicative of a highly refractory mantle source and/or high degrees of

melting beneath arcs, and scandium may be more compatible than sodium during mantle melting. However, it seems to us that either heavy REE, yttrium, titanium, and vanadium are poor indicators of the extent of mantle melting in the source of primitive lavas, or Na2O is a poor indicator, or neither is a good proxy for the extent of mantle melting. Positive correlation between FeO, scandium, vanadium, and titanium, and negative correlations of each of these with SiO2, for all primitive arc lavas raises, once again, the question of whether the concentrations of all these elements are related to FeTi oxide fractionation. However, the plot of TiO2 versus Mg# shows that

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Figure 7 Wt.% Na2O, wt.% TiO2, and ppm Sc versus other trace-element concentrations in ppm, for primitive arc lavas. Again, these plots clearly separate andesites from basalts. Primitive andesites from the oceanic, western Aleutians plot with other primitive andesites from continental arcs. Combined heavy REE, Y, Sc, Ti, and V depletion are indicative of either an important role for garnet fractionation, or a more depleted mantle source, in the genesis of primitive andesites as compared to primitive basalts. Data and symbols as for Figure 6.

Characteristics of Arc Magmas concentration of titanium increases with decreasing Mg# in primitive lavas (Figure 1), in accord with experimental studies showing that primitive lavas at oxygen fugacities typical for arcs are not saturated in FeTi oxides. In addition, heavy REE concentrations are correlated with TiO2 in primitive arc lavas. Heavy REE are incompatible in FeTi oxides and silicates other than garnet (e.g., EarthRef database at http://earthref.org/ and references therein). These two observations almost certainly rule out an important role for FeTi oxide fractionation in controlling the major-element compositions of primitive lavas. Instead, the similar behavior of vanadium, titanium and heavy REE suggests either high degrees of mantle melting, or an important role for garnet fractionation, in the genesis of primitive andesites (see Section 3.18.3.2.5).

3.18.3.2.1

Primitive basalts predominate

Primitive arc basalts have trace-element characteristics that are very distinct from primitive MORB (Figure 3). Figure 4 shows that primitive basalts are more commonly sampled than any other primitive magma type in most oceanic arcs. Biased sampling of picturesque and dangerous strato-volcanoes composed of calcalkaline andesite, rather than low lying basaltic shields, may have complicated this picture for continental arcs, but basalts may predominate among primitive lavas in continental arcs as well (e.g., White and McBirney, 1978). The geochemical characteristics of primitive basalts have been the subject of numerous recent reviews (Elliott, 2003; Davidson, 1996; Elliott et al., 1997; Hawkesworth et al., 1993a,b, 1997; Pearce and Peate, 1995; Plank, 2003; Plank and Langmuir, 1988, 1993, 1998; Tatsumi and Eggins, 1995; Turner et al., 2003, 2001). Considerable uncertainty remains regarding the relative importance of various processes in producing primitive arc basalts. For example, in some arcs the presence of nearly anhydrous, primitive basalts suggests a large role for decompression melting (Bartels et al., 1991; Draper and Johnston, 1992; Elkins Tanton et al., 2001; Sisson and Bronto, 1998), although see also Turner and Foden (2001), with major and trace-element systematics that might be similar to those beneath mid-ocean ridges (e.g., Plank and Langmuir, 1988, 1993). Possible mechanisms for decompression melting include near vertical, diapiric upwelling of low density mixtures of melt þ mantle peridotite (e.g., Davies and Stevenson, 1992; Iwamori, 1997) or diagonal upwelling of peridotite in return flow due to viscous entrainment of the mantle wedge with the subducting plate (Conder et al., 2002; Furukawa, 1993a,b;

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Kelemen et al., 2003a; Kincaid and Sacks, 1997; van Keken et al., 2002). Other evidence implies that decompression may be minor or absent, and “fluxed melting” of the mantle is controlled mainly by addition of H2O and other fluxes from subducting material into the mantle wedge. This has generally been modeled as “fluid-fluxed melting,” resulting from addition of an aqueous fluid to initially solid, but hot peridotite (Abe et al., 1998; Eiler et al., 2000; Grove et al., 2001, 2003; Ozawa, 2001; Ozawa and Shimizu, 1995; Stolper and Newman, 1992). Thus, static or even descending mantle peridotite in the wedge could partially melt if sufficient aqueous fluid were added, provided aqueous fluid reached parts of the mantle where temperature exceeded the fluid-saturated peridotite solidus. Substantial H2O contents in primitive arc basalts, commonly ,3 wt.% (Anderson, 1974; Baker and Eggler, 1987; Falloon and Danyushevsky, 2000; Kamenetsky et al., 1997; Kelemen et al., 1990b; Macdonald et al., 2000; Mu¨ntener et al., 2001; Pichavant et al., 2002; Roggensack et al., 1997; Sisson and Grove, 1993b; Sisson and Layne, 1993; Sobolev and Chaussidon, 1996), but more than 4.5 wt.% in primitive andesites (Grove et al., 2003) are often taken as evidence for addition of H2O-rich fluid to the mantle wedge, and for fluxed melting. A variant on the flux melting hypothesis is the idea of “melt-fluxed melting,” in which reaction between hydrous partial melts of subducting sediment and/or basalt and overlying mantle peridotite leads to increasing melt mass, producing a hybrid “primary melt” in which more than 90% of the compatible elements, such as magnesium, iron, and nickel, are derived from the mantle, while most of the alkalis and other incompatible trace elements come from the initial, melt of subducted material (e.g., Kelemen, 1986, 1990, 1995; Kelemen et al., 1993, 2003b; Myers et al., 1985; Yogodzinski et al., 1995; Yogodzinski and Kelemen, 1998). If this process occurs, it would be facilitated by the fact that melts migrating upwards in the mantle wedge must heat as they decompress; under such circumstances, even anhydrous melts will be able to dissolve solid mantle minerals comprising tens of percent of the initial liquid mass (Kelemen, 1986, 1990, 1995; Kelemen et al., 1993). Similarly, Grove and co-authors propose that fluid-saturated partial melts of peridotite form at ,950 8C, very close to the subducting plate, and then these fluid-saturated peridotite melts cause melt-fluxed melting higher in the wedge (Grove et al., 2003). Aqueous-fluid-saturated melts of eclogite facies sediment or basalt, and fluid-saturated melts of mantle peridotite, would have 25 – 50 wt.% H2O at 3– 5 GPa (Dixon and Stolper, 1995; Dixon et al., 1995; Kawamoto and Holloway, 1997;

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Mysen and Wheeler, 2000). In this regard, they would be efficient fluxing agents, causing additional melt to form via melt/rock reaction in the mantle wedge (Eiler et al., 2000). In principle, addition of such H2O-rich melts to the mantle wedge could explain the substantial water contents in primitive arc magmas, without additional H2O from a fluid. It is worth noting that melt-fluxed melting is distinct from most open-system processes proposed to explain melting beneath mid-ocean ridges, including batch melting (e.g., Gast, 1968; Presnall and Hoover, 1984; Shaw, 1970), fractional melting (e.g., Gast, 1968; Johnson et al., 1990; Langmuir et al., 1977; Richardson and McKenzie, 1994; Shaw, 1970), incremental melting (e.g., Johnson et al., 1990; Kinzler and Grove, 1992; Kinzler and Grove, 1993; Klein and Langmuir, 1987; Langmuir et al., 1977), and continuous melting (Iwamori, 1994; Johnson and Dick, 1992; Langmuir et al., 1977; Sobolev and Shimizu, 1992), because in the latter processes melt forms due to decompression, not as a result of reaction between solid phases and migrating melt. In zone refining (e.g., Harris, 1957; Langmuir et al., 1977), melt mass is constant, so this too is different from melt-fluxed melting. Models of melt generation beneath mid-ocean ridges that include increasing melt mass due to reactive porous flow (e.g., Asimow and Stolper, 1999; Iwamori, 1994; Jull et al., 2002; Lundstrom et al., 1995, 2000; Spiegelman and Elliot, 1992; Spiegelman et al., 2001) are similar to proposed processes of melt-fluxed melting beneath arcs. (Note that Langmuir et al., 1977 mentioned but did not model this process for MORB genesis.) However, beneath arcs— unlike ridges—melt-fluxed melting may be extensive, even in regions that are not simultaneously undergoing decompression melting. Understanding the different mantle melting processes, and determining their relative importance in the generation of primary arc basalts, is an active area of research, with much sponsorship from the US National Science Foundation’s MARGINS Initiative. We anticipate rapid developments in increasingly refined theories on this topic. Another very active area of recent research is the identification of several different source components in primary arc basalts, including (i) fluids derived by dehydration of subducting metabasalt, (ii) fluids derived by dehydration of subducting metasediment, (iii) partial melts of subducting basalt, (iv) partial melts of subducting sediment, (v) fertile mantle peridotite similar to the MORB source, (vi) mantle peridotite depleted by melt extraction beneath a mid-ocean ridge and/or a back-arc basin, and (vii) enriched mantle similar to

the source of ocean island basalt. We return to this topic in Section 3.18.3.3. Nonetheless, as we show in previous and subsequent sections of this chapter, focusing exclusively on arc basalts risks missing endmembers whose characteristics epitomize the difference between arc versus mid-ocean ridge magmas. Thus, in the following sections we focus on other types of primitive arc lavas. However, some detailed characteristics of primitive arc basalts, together with other primitive arc magmas, are described in Section 3.18.3.3. 3.18.3.2.2

Are some low Mg# basalts primary melts? perhaps not

Hypothetical derivation of primary, low Mg# basalts and andesites from partial melting of subducted basalt—without major-element equilibration with the overlying mantle—remains controversial for arcs (e.g., Brophy and Marsh, 1986; Johnston and Wyllie, 1988; Marsh, 1976; Myers et al., 1986a,b, 1985) as well as hotspots (e.g., Chauvel and Hemond, 2000; Hauri, 1995; Korenaga and Kelemen, 2000; Lassiter and Hauri, 1998; Sobolev et al., 2000) and even mid-ocean ridges (e.g., Schiano et al., 1997). Based on the criteria outlined by Gill (1981, 1974, 1978), we see no evidence for direct partial melts of subducted, eclogite facies sediment or basalt in our data compilation. It may be that diapirs of melting basalt always rise to depths at which garnet is no longer stable, prior to separation of melt from residue (Brophy and Marsh, 1986), but there seems to be no direct evidence for this. Instead, arc lavas with a trace-element signature consistent with derivation via partial melting of eclogite (e.g., high middle/heavy REE ratios) are primitive, with Mg# . 60 (e.g., Grove et al., 2001; Kay, 1978; Kelemen et al., 2003b; Yogodzinski et al., 1995; Yogodzinski and Kelemen, 1998). Thus, in this chapter we make the simplifying assumptions that lavas with Mg# , 60 are derived from primitive melts via crustal differentiation, and that all melts passing from the mantle wedge into arc crust have Mg# . 65 (depending on Fe3þ/Fe2þ), and are close to Fe/Mg exchange equilibrium with mantle peridotite (Mg# , 70). Related to this topic are questions about the genesis of tonalites, trondhjemites, and granodiorites (TTGs), that are common in Archean cratons (see Chapter 3.11). Although it is difficult to be certain, we believe that TTGs are probably not primary melts of subducting eclogite as has been proposed (Defant and Kepezhinskas, 2001; Martin, 1986, 1999; Rapp and Watson, 1995; Rapp et al., 1991), simply because it does not seem likely that H2O-rich, low temperature melts could traverse the high temperature mantle wedge without substantial

Characteristics of Arc Magmas reaction with peridotite. Instead, we infer that TTGs are probably the products of intracrustal differentiation, with felsic melts rising to the upper crust, and mafic residues remaining in the lower crust. (Note that, while seismic and petrologic data on continental crust clearly establish that it is differentiated, the intracrustal differentiation process could have modified an initially andesitic or basaltic bulk composition.) They may have evolved by crystal fractionation from a parental, primitive andesite melt. However, because the process of intracrustal differentiation may have involved residual garnet, it is difficult to discern which TTGs with heavy REE depletion inherited their trace-element characteristics from primitive andesites, and which reflect crustal garnet fractionation. 3.18.3.2.3

Boninites, briefly

In determining the characteristics of “primary” arc magmas—melts that pass from residual mantle into the overlying, igneous crust—most recent reviews of arc geochemistry have concentrated on the characteristics of primitive basaltic magmas. In doing so, these reviews have implicitly incorporated the assumption that primary arc magmas are invariably basaltic. In our view, there are two types of “andesitic” primitive magmas in arcs, boninites and primitive andesites. These two types of magmas extend to end-members having Mg# . 0.7, and carry olivine phenocrysts with Mg# . 90 (typical mantle values). While these are less common than primitive basalts, we think they are important for the reasons enumerated in Section 3.18.3.1.3. While there are many far more detailed definitions and subdivisions of boninite lava compositions (e.g., Crawford, 1989), we found it convenient to simply define boninites as lavas with .54 wt.% SiO2, ,0.5 wt.% TiO2 ( plus samples described in the original data sources as boninites, including some basalts). As can be seen in Figures 1, 6, and 7, lavas defined in this way share many other distinctive characteristics, including high MgO and low alkali contents at a given SiO2 and Mg#. Boninites are largely restricted to western Pacific island arcs, and in those arcs they are apparently more abundant in the early stages of magmatism (e.g., Bloomer and Hawkins, 1987; Falloon et al., 1989; Stern and Bloomer, 1992). Their high MgO contents (some .10 wt.%) and the presence of clinoenstatite phenocrysts, probably reflect both high temperatures and high water contents in the mantle wedge, with a highly depleted, harzburgite residue, consistent with generally low REE concentrations and flat to light REE depleted patterns (e.g., Falloon and Green, 1986; Falloon

613

et al., 1989; Pearce et al., 1992; Sobolev and Danyushevsky, 1994). Most authors accept that most boninites are derived by crystal fractionation from primary andesite melts derived by high degrees of relatively low pressure melting, with a harzburgite residue. However, some lavas termed boninites could conceivably be derived via substantial olivine ^ low calcium pyroxene fractionation from very high Mg#, incompatible-element-depleted, primary picrites or komatiites. The abundance of boninites in early stages of western Pacific arc magmatism, combined with the high magmatic fluxes inferred for the early stages of magmatism in those arcs, may have led to bulk crustal compositions that remain dominantly boninitic. Relatively low bulk crustal seismic velocities in the Izu – Bonin arc (Suyehiro et al., 1996), compared to the central Aleutian arc (e.g., Holbrook et al., 1999), might reflect higher SiO2 in the Izu –Bonin crust. Nonetheless, boninitic crust with .54 wt.% SiO2 would be depleted in alkalis and light REEs, and thus very different from continental crust, and from calc-alkaline magma series.

3.18.3.2.4

Primitive andesites: a select group

We turn now to primitive andesites (Mg# . 0.6) and high Mg# andesites (Mg# . 0.5). High Mg# andesites have been the subject of much attention in recent years because of their unique major and trace-element characteristics (e.g., Baker et al., 1994; Defant et al., 1991a; Defant and Drummond, 1990; Defant et al., 1992, 1989, 1991b; Defant and Kepezhinskas, 2001; Grove et al., 2001, 2003; Kay, 1978; Kelemen, 1995; Kelemen et al., 2003b; Rogers et al., 1985; Shimoda et al., 1998; Stern and Kilian, 1996; Tatsumi, 1981, 1982, 2001a; Tatsumi and Ishizaka, 1981, 1982; Yogodzinski and Kelemen, 1998, 2000; Yogodzinski et al., 1995, 2001, 1994). Some high Mg# andesites—particularly some lavas on Adak Island in the Aleutians (Kay, 1978)—have been called “adakites” (e.g., Defant and Drummond, 1990) as well as “sanukitoids,” “high Mg andesites,” and “bajaites.” The term “adakite” is used in a variety of contexts by different investigators, but generally refers to andesites and dacites with extreme light REE enrichment (e.g., La/Yb . 9), very high Sr/Y ratios (e.g., Sr/Y . 50), and low yttrium and heavy REE concentrations (e.g., Y , 20 ppm, Yb , 2 ppm). In the Aleutians, all lavas with these characteristics are high Mg# andesites and dacites. However, the de facto definition of “adakite” does not specify a range of Mg#. Worldwide, many evolved lavas have been termed adakites. Thus, not all adakites are high Mg#

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Geochemistry of Subduction-related Magmatic Arcs

andesites. Similarly, most high Mg# andesites, in the Aleutians and worldwide, have La/Yb , 9 and Sr/Y , 50, and so not all high Mg# andesites are adakites. Finally, for some authors adakite has a genetic connotation. Some investigators infer that all andesites and dacites with extreme light REE enrichment, very high Sr/Y ratios, and low yttrium and heavy REE concentrations formed via partial melting of subducted basalt in eclogite facies, and use the term adakite to refer to both composition and genesis interchangeably. While we believe that many high Mg# andesites do indeed include a component derived from partial melting of eclogite, we feel it is important to separate rock names, based on composition, from genetic interpretations. For this reason, we do not use the term adakite in this chapter. 3.18.3.2.5

Three recipes for primitive andesite

Loosely speaking, the difference between primitive basalts and primitive andesites might arise in several ways. (i) Both may arise from melting of different sources, with primitive andesites incorporating a relatively large proportion of melts of subducted basalt and/or sediment, compared to primitive basalts. (ii) They might arise from the same mantle source, with different degrees of melting, related to different extents of enrichment via fluids derived from subducting sediment and/or oceanic crust. (iii) High Mg# andesites might arise via mixing of primitive basalts with evolved, high SiO2 melts, or assimilation of “granitic” rocks in primitive basalts. We briefly expand on each of these in the next few paragraphs. (i) Primary andesite magma with an eclogite melt component. High Mg# andesites may incorporate a component formed by partial melting of subducted basalt or sediment in eclogite facies, which subsequently reacted with the overlying mantle peridotite to form a hybrid melt (Carroll and Wyllie, 1989; Kay, 1978; Kelemen, 1986, 1995; Kelemen et al., 1993, 2003b; Myers et al., 1985; Yogodzinski et al., 1995, 1994). In the hybrid melt, high incompatible-element contents reflect eclogite melting, and major-element concentrations reflect melt/mantle equilibration. In this view, high H2O, K2O, and Na2O contents stabilize high SiO2 melt in equilibrium with mantle olivine at ,1 GPa, as demonstrated experimentally for simple systems (Hirschmann et al., 1998; Kushiro, 1975; Ryerson, 1985), peridotite melting experiments (Hirose, 1997; Kushiro, 1990; Ulmer, 2001), and phase equilibrium experiments on natural primitive andesite compositions (e.g., Baker et al., 1994; Grove et al., 2003; Tatsumi, 1981, 1982).

In some cases the entire incompatible traceelement budget of these hybrid melts might be derived from eclogite melting, with only major elements and compatible trace elements (nickel, chromium) affected by interaction with peridotite. However, there are few primitive lavas in our compilation with a clear eclogite melting signature. Since heavy REEs and yttrium are compatible in garnet, heavy REE and yttrium concentrations in eclogite melts should be low, and middle to heavy REE ratios in eclogite melts should be high (e.g., chondrite normalized Dy/Yb . 1.5). The few primitive lavas in our compilation that do have chondrite normalized Dy/Yb . 1.5 are mainly primitive andesites from the western Aleutian arc, at and west of Adak Island. Thus (Figure 8), not all light REE enriched, high Sr/Nd arc lavas have high Dy/Yb. Lack of a clear eclogite melting signature in heavy REE and yttrium contents does not rule out a role for eclogite melt in producing high Mg# andesites. In hypothesis (i), the concentrations of heavy REE and yttrium in primitive arc andesites are interpreted as having been raised by reaction of eclogite melt with mantle peridotite at moderate melt/rock ratios (Kelemen, 1995; Kelemen et al., 1993, 2003b). Modeling shows that this process can produce a very close match to most high Mg# andesite compositions (e.g., Kelemen et al., 2003b; figures 21D and 21E). In this interpretation, primitive arc basalts incorporate a finite but smaller amount of an eclogite melt component (e.g., Kelemen et al., 2003b; figure 21F). As far as we know, there are few if any petrological or geochemical arguments that can be used to rule out this hypothesis. In fact, excluding heavy REEs and yttrium, the incompatible trace-element abundances in the “subduction component” (McCullouch and Gamble, 1991) inferred from inversion of major and trace elements in Marianas back-arc lavas (Stolper and Newman, 1992), an array of western Pacific arc lavas focused on the Vanuatu arc (Eiler et al., 2000), and primitive lavas from Shasta volcano in the southern Cascades (Grove et al., 2001) closely resemble trace-element concentrations experimental and predicted partial melts of eclogite (Kelemen, 1995; Kelemen et al., 1993, 2003b; Rapp et al., 1999), and of erupted high Mg# andesites. One inverse model that is consistent with all available major-element, trace-element and isotopic constraints (Eiler et al., 2000, p. 247) involves reaction between mantle peridotite and a silicate melt derived from subducting eclogite, with 30 wt.% H2O (appropriate for fluid-saturation at 3– 5 GPa, Dixon and Stolper (1995), Dixon et al. (1995), Kawamoto and Holloway (1997), Mysen and Wheeler (2000)).

Characteristics of Arc Magmas 4.0

4.0

2.5 2.0 1.5

tl e

3.0 2.5

0.5

0.5 Plag fr'n 100

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nw

/m

1.5 1.0

3

ti o R e ac

2.0

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0.0

an

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Dy/Yb (N)

Reaction w/ mantle

3.5

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0.0 1

10 La/Yb (N)

Figure 8 Relationship between Sr/Nd, La/Yb, and Dy/Yb for primitive arc lavas. High Dy/Yb is probably indicative of an important role for residual garnet in the genesis of some lavas. Not all lavas with very high La/Yb and Sr/Nd have high Dy/Yb. Thus, it may be unwise to use La/Yb and high Sr as indications that a given igneous rock is derived from a source with abundant, residual garnet. Arrows labeled “reaction with mantle” show results for traceelement models of reaction of a partial melt of MORB in eclogite facies with upper mantle peridotite (Kelemen et al., 2003b). Arrow marked “plag fr 0 n” emphasizes that, because Sr is much more compatible than Nd in plagioclase, crystal fractionation of plagioclase, or crystal assemblages with cotectic proportions of plagioclase, leads to decreasing Sr/Nd. Data and symbols as for Figure 6. Sr/Nd on x-axis in left panel is not normalized, and is on a logarithmic scale.

Nevertheless, this hypothesis has been unpopular since the 1980s, because geodynamic models, incorporating either constant mantle viscosity or a rigid upper plate of prescribed thickness, predicted that solidus temperatures could not be reached in basalt or sediment at the top of the subducting plate, except under unusual circumstances (see reviews in Kelemen et al., 2003a; Peacock, 1996, 2003; Peacock et al., 1994). For example (Eiler et al., 2000, p. 247), discounted their successful model involving H2O-rich silicate melt because “[the] successful melt-fluxed [model]…require[s]…temperatures…that are…not obviously compatible with…thermal models… (Peacock, 1996).” Recently, as discussed in Section 3.18.3.3.4, subduction zone thermal models that incorporate thermally dependent viscosity and/or non-Newtonian viscosity in the mantle wedge predict temperatures higher than the fluid-saturated solidus near the top of the subducting plate beneath arcs at normal subduction rates and subducting plate ages (Kelemen et al., 2003a; van Keken et al., 2002). While this is an area of active research, it is no longer the case that thermal models “rule out” partial melting of subducted material in eclogite facies. (ii) Primary andesite magma from fluxed melting. A more popular model for arc magma genesis is “fluid-fluxed melting” (Section 3.18.3.2.1). An aqueous fluid derived from subducted basalt and sediment enriches the mantle source of arc magmas

in “mobile elements,” while simultaneously causing partial melting of that source (e.g., Abe et al., 1998; Eiler et al., 2000; Grove et al., 2001, 2003; Ozawa, 2001; Ozawa and Shimizu, 1995; Stolper and Newman, 1992; Vernieres et al., 1997). In this interpretation, high LILEs and light REEs are seen as the result of fluid enrichment, while low heavy REEs, titanium, and scandium are seen as the results of high degrees of melting. Again, high H2O, K2O, and Na2O contents derived from the fluxing fluid may lead to relatively high SiO2 melt in equilibrium with peridotite. Following this reasoning, primitive andesites could be the extreme products of fluxed melting (Grove et al., 2003). A series of recent papers describes how this process can produce a very close match to high Mg# andesites from Mt. Shasta volcano in the southern Cascades (Grove et al., 2001, 2003). However, as in previous inversions based on a variety of arc lavas (Eiler et al., 2000; Stolper and Newman, 1992), the required fluids have dissolved light REEs and thorium contents larger than predicted from experimental fluid/rock partitioning studies (Ayers et al., 1997; Brenan et al., 1996, 1995a,b; Kogiso et al., 1997; Stalder et al., 1998; Tatsumi and Kogiso, 1997). If the fluxing agent in flux melting were a silicate melt rather than an aqueous fluid, predicted REE and LILE contents would be much higher, potentially resolving this discrepancy. We return to this point in Section 3.18.3.3.1. From this perspective,

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Geochemistry of Subduction-related Magmatic Arcs

“melt-fluxed melting” can be considered more or less identical to (i). (iii) Mixing of primary basalt and granitic lower crustal melts. Although there are significant fluid mechanical barriers to such a process (e.g., Campbell and Turner, 1985), many chemical features of high Mg# andesites could be explained as the result of mixing of primitive arc basalt with evolved, silica-rich melt with high LILEs and light REEs, and low heavy REEs, titanium, and scandium. In the western Aleutians, high Mg# andesites have abundant, zoned phenocrysts which probably do reflect magma mixing processes. However, the most light REE enriched, heavy REE depleted magmas have the highest Mg# (Kelemen et al., 2003b; Yogodzinski and Kelemen, 1998), which is inconsistent with the hypothesis of mixing primitive basalt with enriched granitic melt outlined in the previous paragraph (Kelemen et al., 2003b; tables 18 and 19). Instead, mixing apparently combined primitive, light REE enriched andesites with more evolved, less enriched andesites. With the exception of the western Aleutian arc, primitive and high Mg# andesites (excluding boninites) are rare in intra-oceanic arcs (Figure 4). High Mg# andesites are most common in continental arcs, where interaction between basalt and pre-existing crust might be important. Primitive andesites have been reported from the Cascades (Baker et al., 1994; Grove et al., 2001, 2003; Hughes and Taylor, 1986), Baja California (Rogers et al., 1985), southeast Costa Rica and western Panama (de Boer et al., 1988, 1995; Defant et al., 1991a,b, 1992), Ecuador (Beate et al.,

2001; Bourdon et al., 2002; Monzier et al., 1997), Argentina (R. W. Kay and S. M. Kay, 1991, 1993), southern Chile (Sigmarsson et al., 2002, 1998; Stern and Kilian, 1996), the Philippines (Defant et al., 1989; Maury et al., 1992; Schiano et al., 1995), Papua New Guinea ((Arculus et al., 1983), SW Japan (Shimoda et al., 1998; Tatsumi, 1982, 2001a,b; Tatsumi and Ishizaka, 1981, 1982), and Kamchatka (Kepezhinskas et al., 1997). In all of these localities, other than the Aleutians, most high Mg# andesites have elevated 208 Pb/204Pb, compared to MORB (Figure 9). Thus, lead isotope data suggest the presence of a component derived either from recycling of lead from subducting sediment, or from crustal interaction of primitive basalts with older, continental crust and continentally derived sediment. Given the fact that so many primitive andesites are in “continental” arcs, crustal interaction processes must be considered. That said, primitive andesites and basalts have overlapping 87Sr/ 86 Sr and 143 Nd/144Nd, which restricts the range of crustal sources that could be involved in mixing or assimilation to create primitive andesites from basalts. Lower crustal anatexis in arcs like the Andes (e.g., Babeyko et al., 2002), that have thick crust and probably garnet at the base of the crust, might be expected to yield appropriate mixing end-members. Indeed, our compilation includes dacitic to rhyolitic lavas from the Andes with high Dy/Yb (Bourdon et al., 2000; S. M. Kay and R. W. Kay, 1994; Matteini et al., 2002). Mixtures of these compositions with primitive basalt have most of the major and trace-element characteristics of high Mg# andesites (e.g., Figure 10).

40.5 0.714 40.0 0.712 39.5

87Sr/86Sr

208Pb/204Pb

0.710 39.0 38.5

0.708

38.0

0.706

37.5

0.704

37.0 17.0

17.5

18.0

18.5 19.0 206Pb/204Pb

19.5

20.0

20.5

0.702 0.5120

0.5125

0.5130

143Nd/144Nd

Figure 9 Pb, Sr, and Nd isotopes in primitive lavas from our arc compilation. In general, primitive andesites do not have distinctive isotopic characteristics, compared to primitive basalts. Primitive andesites from the western Aleutians have the most depleted values in our data compilation. Data and symbols as for Figures 1 and 6.

Characteristics of Arc Magmas

617

4.0 3.5

Garnet granulite melt (gg) mixing w/ primitive basalt (pb)

Dy/Yb (N)

3.0

Hybrid eclogite/peridotite melt (he) reacting w/ spinel peridotite to form primitive andesite (pa) and basalt

gg

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he

2.0 1.5 xl'n

1.0

xl'n (pb) pa

pb

0.5 0.0 0

10

20

30

40 50 Molar Mg#

60

70

80

4.0 3.5

Garnet granulite melt (gg) mixing w/ primitive basalt (pb)

3.0

gg

Dy/Yb (N)

2.5

Hybrid eclogite/peridotite melt (he) reacting w/ spinel peridotite to form primitive andesite (pa) and basalt (pb)

?

he

2.0 1.5 xl'n 1.0

pb

?

xl'n (pb) pa

0.5 0.0 1

10

100 Sr/Nd

Figure 10 Relationship between Mg#, Dy/Yb, and Sr/Nd in arc lavas from our compilation. Partial melts of subducted, eclogite facies MORB have high Dy/Yb and Sr/Nd (e.g., Gill, 1974, 1978; Kay, 1978; Kelemen et al., 1993, 2003b; Rapp et al., 1999; Yogodzinski and Kelemen, 1998). When these melts react with overlying mantle, the hybrid liquid acquires high Mg# (Kay, 1978; Kelemen et al., 1993, 2003b; Rapp et al., 1999; Yogodzinski and Kelemen, 1998). Western Aleutian primitive andesites may form in this way. Continued reaction increases heavy REE and Y contents to values in equilibrium with depleted mantle peridotite as melt/rock ratios decrease to ,0.1 or less (Kelemen et al., 1993, 2003b). High Sr/Nd, primitive andesites and even primitive basalts Dy/Yb , 1.5 may form in this way. Alternatively, primitive andesites could be mixtures of lower crustal melts and primitive basalt (e.g., S. M. Kay and R. W. Kay, 1994). Partial melts of lower crustal garnet granulites, as exemplified by some Andean andesite and dacites (Bourdon et al., 2000; S. M. Kay and R. W. Kay, 1994; Matteini et al., 2002), have high Dy/Yb at lower Mg# and Sr/Nd than Aleutian primitive andesites. Mixing with primitive basalt (e.g., average Marianas primitive basalt, Table 1) yields high Dy/Yb, high Mg# melts similar to Aleutian primitive andesites. However, such mixtures have low Sr/Nd, unlike Aleutian and other primitive andesites. Because Sr is more compatible than Nd in plagioclase, this result is likely to be general. Thus, while magma mixing may have played some role in the genesis of many or most arc lavas, mixing is an unlikely explanation for the genesis of primitive andesite compositions.

Thus, in this chapter we re-examine the hypothesis that primitive andesites are produced by crustal mixing or assimilation using global data, rather than simply data from the Aleutians. This reexamination could be extended to the western Aleutians as well; perhaps, small proportions of high SiO2, small-degree melts of garnet granulite in arc lower crust are commonly produced but

rarely erupted. Could such mixing explain the composition of primitive andesites in general? Key data to address this question come from ratios of trace elements that are fractionated differently by partial melting of eclogite compared to garnet granulite. For example, because strontium is compatible in plagioclase, strontium is more incompatible than neodymium in melting of

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Geochemistry of Subduction-related Magmatic Arcs

plagioclase-free eclogite, whereas strontium is more compatible than neodymium in partial melting of lower crustal granulite with substantial plagioclase. Because temperatures at the base of the crust in active arcs are high (.800 8C) and crustal thickness rarely exceeds 60 km, eclogite facies assemblages will not form in arc lower crust. However, formation of garnet granulites is possible where crustal thickness exceeds 25 – 30 km (e.g., Jull and Kelemen, 2001). Although garnet formation might be inhibited by slow kinetics in H2O-poor lower crustal cumulates, observations from the Talkeetna and Kohistan arc crustal sections show that garnet did form at pressures where it was thermodynamically stable (Section 3.18.4). While some garnet in the Talkeetna lower crustal section is entirely metamorphic in origin, other samples include residual or cumulate, igneous garnet (Section 3.18.4.3). As noted above, our arc lava compilation includes high Dy/Yb, evolved lavas from the Andes, which are probably partial melts of garnet granulite. However, while these lavas are appropriate in other ways as mixing end-members to produce high Mg# andesites, they have low Sr/Nd (Figure 10). Therefore, these evolved Andean lavas could not mix with primitive arc basalt to produce high Mg# andesites, which have higher Sr/Nd (Figure 10), and it is unlikely that any other lower crustal melts would be appropriate. In many ways, the process of “assimilation,” or melt/rock reaction, is comparable to magma mixing. Reaction between hot, primitive basalt and granitic wall rock, in particular, shares many characteristics with mixing of primitive basalt and a granitic partial melt of wall rock. For example, both processes tend to produce high, nearly constant compatible-element concentrations and ratios (magnesium, iron, nickel, Mg#) together with substantial enrichment in incompatible trace elements, over much of the range of mixing or reaction progress (compare, e.g., DePaolo, 1981; and Kelemen, 1986 with O’Hara and Mathews, 1981). In specific cases the outcome of melt/rock reaction may be quite distinct from magma mixing. Selective dissolution of plagioclase in hydrous, plagioclase-undersaturated melt could enrich resulting liquids in Sr/Nd, potentially producing trends distinct from the mixing trends in Figure 10. However, note that selective dissolution of plagioclase, alone, might create a telltale anomaly with high Eu/Sm, and would not explain other characteristics of high Mg# andesites, including heavy REE, titanium, scandium, and vanadium depletion. Instead, selective dissolution of plagioclase would have to be coupled with crystallization of garnet in order to explain high Mg# andesite genesis via

crustal melt/rock reaction. In view of the fact that garnet is not saturated in primitive arc melts at pressures of ,1 –1.5 GPa (e.g., Mu¨ntener et al., 2001), we view this as unlikely. Furthermore, on a global basis, high Sr/Nd is negatively correlated with 87Sr/86Sr, ruling out assimilation of continental granitoids, or sediments derived from continental crust. (iv) Primitive andesite summary. In summary, primitive andesites are probably derived from primary andesite magmas, produced by processes below the base of igneous arc crust, which are different from primary magmas of primitive arc basalts. Mixing of primitive basalt and evolved partial melts of lower crustal garnet granulite probably cannot produce end-member high Mg# andesite lava compositions. Similarly, lower crustal assimilation probably cannot produce typical high Mg# andesites. Instead, we conclude that most primitive andesites are probably produced via process (i), described above, by reaction of small degree partial melts of subducted, eclogite-facies sediment and/or basalt with the overlying mantle wedge. In our view, process (ii), melting of the mantle fluxed by an enriched “fluid” component derived from subducted sediment and/or basalt, is only viable if the fluxing “fluid” is, in fact, a melt or a supercritical fluid with partitioning behavior similar to melt/rock partitioning. Thus, viable versions of process (ii) are the same as process (i). Primary magmas parental to primitive andesites may also be parental to calc-alkaline, evolved arc magmas. It is evident that fractionation of olivine, pyroxene, and/or plagioclase from a primitive andesite melt leads to higher SiO2 and alkali contents, at a given degree of crystallization, compared to fractionation of the same phases from primitive basalt (Figure 11). The key here is that olivine þ pyroxene cumulates, and olivine þ pyroxene þ plagioclase cumulates, have ,50% SiO2, as determined by mineral stoichiometry and cotectic proportions. Crystallizing these solid assemblages from a basalt with ,50% SiO2 does not change the SiO2 content of the resulting liquid, whereas removing the same cumulates from an andesite with 55 wt.% or even 60 wt.% SiO2 leads to an increase in SiO2 in the derivative liquid. (Note that the variation in Figure 11, taken out of context, might also be attributed to fractionation from a common parental magma with Mg# ,0.8; however, trace-element and isotope variation precludes this possibility.) The high SiO2 and—probably—H 2O contents of primitive andesites may make them difficult to erupt. As they reach the mid-crust, become saturated in H2O, and degas, their viscosity must rise abruptly, leading to slower melt transport and enhanced rates of crystallization. For these reasons,

Characteristics of Arc Magmas

619

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80 70

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Figure 11 On the left, a schematic illustration showing how fractionation of primitive pyroxenite or gabbro from primitive basalt leads to decreasing Mg# at nearly constant SiO2, while fractionation of the same crystal assemblages from primitive andesite leads to increasing SiO2. On the right, data from the oceanic Aleutian arc (compiled by Kelemen et al., 2003b) show how variation between lava series from different volcanoes might arise as a result of this effect. Sources of data for Seguam, Okmok and Piip volcanoes are given in the caption for Figure 2. Filled symbols for each volcano were used in power law curve fits. Open symbols were omitted.

primitive andesites may be more common among melts entering the base of the igneous crust than they are among erupted lavas. It is not clear to what extent this analysis can be extended to high Mg# andesite compositions typical of plutonic rocks in calc-alkaline batholiths. Some Aleutian and Cascades plutonic rocks are high Sr/Nd, high Mg# andesites which cannot be produced via crustal mixing, but high Sr/Nd in these rocks could arise via incorporation of cumulate plagioclase. Many calcalkaline plutonic rocks have relatively low Sr/Nd, and could be mixtures of lower crustal melts and primitive basalt. Because plagioclase crystallization leads to decreasing Sr/Nd with decreasing Mg#, however, most calc-alkaline plutonic rocks could also be derived via crystal fractionation from primitive andesite. Therefore, the Sr/Nd discriminant between lower crustal melts (low Sr/Nd) and eclogite melts (high Sr/Nd) is only useful for lavas that retain high Dy/Yb. Unfortunately, this means that Sr/Nd cannot be used to determine the extent to which the high Mg# andesite composition of continental crust is due to crystal fractionation from a primitive andesite parent, versus the extent to which it is due to mixing of primitive basalt and lower crustal melts. (v) Why are primitive andesites rare? In Sections 3.18.3.2.5 and 3.18.3.3, we argue that a component derived from partial melting of

subducting sediment and/or basalt is included in most arc magmas. This is consistent with a substantial body of work calling upon partial melts of subducting sediment to explain traceelement enrichments in arc basalts, but inconsistent with the theory that primitive andesites with “adakite” trace-element signatures (Section 3.18.3.2.4), which apparently include tens of percent eclogite melt (Kelemen et al., 2003b), are only found in arcs with unusually hot subduction zones—due to subduction of young oceanic crust, very slow convergence rates allowing substantial time for conductive heating, and/or discontinuous “tears” in the subducting plate which enhance mantle convection and allow conductive heating from the side as well as the top and bottom (e.g., de Boer et al., 1991, 1988; Defant and Drummond, 1990; Yogodzinski et al., 1995, 2001, 1994). If partial melts of subducted material are ubiquitous in arcs, why do they form large proportions of arc magma in a few places, and very small proportions (a few percent, e.g., Class et al. (2000)) in most arcs? Following (Kelemen et al., 2003b), we offer the following tentative explanation. Arcs that have primitive andesites are similar in having slow convergence rates, and it may be true that many are situated above “tears” in the subducting plate. Most (except the Aleutians) are in regions of young plate subduction. Thus, the subduction

620

Geochemistry of Subduction-related Magmatic Arcs 100

ppm La

ppm Ba

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10

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10 ppm Nd

100

Figure 12 Trace-element concentrations, in ppm, for primitive lavas in our compilation. In general, primitive arc lavas are enriched in Th, Ba, La, Pb, Ce, Sr, and Nd, compared to average MORB. (Concentrations of these elements in primitive MORB are even lower than the average values). This is somewhat at odds with theories invoking separate enrichment processes involving aqueous fluids for Ba, Pb, and Sr, and partial melts of subducting sediment for Th, La, Ce, and Nd. Diagonal lines illustrate constant trace-element ratios. The overall trends of the compiled data do not diverge dramatically from constant ratios of these elements (diagonal lines). If anything, “fluid-mobile” Ba, Pb, and Sr show less variation and less enrichment relative to MORB than “immobile” Th and La. Since all of these elements are highly incompatible during melting of eclogite facies sediment and basalt, and in peridotite/melt equilibria,

Characteristics of Arc Magmas zone may be unusually hot, and may yield a larger proportion of partial melt. However, it is hard to imagine that this can account for a factor-of-ten difference in the proportion of subduction zone melt in arc lavas. Instead, other factors may be involved. Slow convergence leads to slow convection in the mantle wedge, enhancing the amount of conductive cooling. This could be very important, because of positive feedback: increased viscosity due to cooling could slow wedge convection still further. Thus, we suggest that primitive andesites are found in areas in which the subducting plate is relatively hot, producing more partial melt, and the overlying wedge is relatively cold, producing less partial melt. In contrast, in normal arcs, abundant melts derived from the mantle wedge overwhelm the subduction zone melt signature, except for distinctive isotope ratios and thorium contents carried in partial melts of sediment. In addition, as outlined in Sections 3.18.3.1.3 and 3.18.3.1.4, primitive andesites may contain more H2O than primitive basalts (e.g., Grove et al., 2003). If so, they will become H2O saturated and degas at mid-crustal depths. As a result of degassing, these magmas will undergo rapid crystal fractionation, so that only evolved andesites and dacites reach the surface. Also, as a result of degassing together with their relatively high SiO2 contents, primitive andesites will become very viscous in the mid-crust, and this may inhibit their eruption.

3.18.3.3 Trace Elements, Isotopes, and Source Components in Primitive Magmas 3.18.3.3.1

Incompatible trace-element enrichment

Lava compositions from arcs worldwide share many characteristics that are, in turn, very distinct from those of MORBs. These distinctions have been known for decades, and we cannot hope to provide a comprehensive review of the entire literature on this topic. However, we do wish to use our data compilation to quantify the differences between arc and MORB lavas. Thus, caesium, rubidium, radium, barium, thorium, uranium, and potassium are enriched in most arc lavas relative

621

to MORB, together with light REEs, potassium, lead, and strontium (Figure 12). Moderately incompatible elements (middle- to heavy-REE and titanium) are generally depleted relative to MORB (Figure 6). These characteristics are present in both oceanic and continental arcs. Globally, these characteristics are observed even in the most primitive lavas, and some of the most enriched lavas in our data set are western Aleutian primitive andesites with low 87Sr/86Sr and lead Yogodzinski et al., 1995, isotope ratios (Kelemen et al., 2003b; 1994). Thus, it seems likely that the enrichments do not arise mainly as a result of crustal processes, and instead are present in melts entering the base of arc crust, and have a subMoho origin, at least in part. Based on similar reasoning, a host of studies in the 1980s and 1990s were designed to decipher the subcrustal source of enriched incompatible traceelement contents in arc magmas. Possible sources of enrichment, relative to the MORB mantle source, include (i) aqueous fluids derived by dehydration of subducting metabasalt, (ii) aqueous fluids derived by dehydration of subducting metasediment, (iii) aqueous fluids derived by dehydration of partially serpentinized mantle peridotite, (iv) hydrous partial melts of subducting basalt, (v) hydrous partial melts of subducting sediment, and (vi) the presence of “enriched mantle” similar to the various mantle source components inferred for ocean island basalt. These could act upon, or mix with (vii) fertile mantle peridotite similar to the MORB source, or (viii) mantle peridotite depleted by melt extraction beneath a mid-ocean ridge and/or a back arc basin. Given the relatively poorly known (and probably variable) compositions of these various components, their poorly known (and variable) proportions in a given primary arc magma, and the poorly known processes through which different components might interact with additional trace-element fractionation, there may be numerous combinations of these components which could account for the composition of a given primary magma. As a result, most authors have tried to simplify the geochemical interpretation of arc petrogenesis by concentrating on just a few of these components, whose interaction might account for much of the trace-element variability in arc basalts.

similar enrichment of all of them is a natural consequence of melting and melt transport. Conversely, because Th, La, Ce, and Nd are less soluble in aqueous fluids than Ba, Pb, and Sr, dehydration and aqueous fluid transport should fractionate these elements. Large gray symbols show estimated compositions of fluid (rectangles) and melt (circles) in equilibrium with eclogite for Marianas (dark gray) and Aleutians (light gray) at 2 wt.% fluid or melt extracted (Table 2(b)). Size of rectangle reflects results obtained using the range of different partition coefficients. Nd concentration is not calculated and is assumed to be the same as Ce concentration (not shown in Table 2). Arrow outline labeled “plag fr0 n” in diagrams involving Sr reminds readers that even small amounts of plagioclase crystallization will lead to decreasing Sr together with increasing concentrations of Th and Nd. Symbols as in Figures 1 and 6.

0.7070

0.7060

0.7060

0.7050

0.7050

0.7040

0.7040

0.7030

0.7030

0.7020

Ba/La

0.7070

87Sr/86Sr

87Sr/86Sr

100.00

0.00

0.7020 0

10

20

30

40

50 60 Ba/La

70

80

90 100

50.00

0.00 0.10 0.20 0.30 0.40 0.50 0.60 0.70 0.80 Th/La

0.00 0.10 0.20 0.30 0.40 0.50 0.60 0.70 0.80 Th/La

0.7070

0.7060

0.7060

0.7050

39.00

208Pb/204Pb

0.7070

87Sr/86Sr

87Sr/86Sr

39.50

0.7050

0.7040

0.7040

0.7030

0.7030

10

100 Ba (ppm)

1,000

38.00

37.50

0.7020

0.7020

38.50

37.00 1.0 Th (ppm)

10.0

1.00 Th (ppm)

10.00

Characteristics of Arc Magmas (i) Three main components? Using traceelement ratios. Particularly influential have been studies which call upon three principal components in arc basalt petrogenesis: (a) aqueous fluids derived from metabasalt, (b) partial melts of subducting sediment, and (c) MORB source mantle. (e.g., Class et al., 2000; Ellam and Hawkesworth, 1988a; Elliott, 2003; Elliott et al., 1997; Hawkesworth et al., 1997; Johnson and Plank, 1999; Miller et al., 1994; Plank and Langmuir, 1993, 1998; Turner et al., 2003, 2000a,c). These studies generally rely on analyses of both primitive and evolved basalts, and so they use incompatible trace-element ratios in an attempt to remove the effects of crystal fractionation. In these studies, one component, with relatively low Ba/La and Th/La, and MORB-like isotope ratios is interpreted as the pre-existing mantle source, prior to fluid and sediment melt enrichment. Also in these studies, enrichments in “fluid mobile” elements relative to light REE (e.g., high U/Th, Ra/Th, Ba/La, Pb/Ce, and Sr/Nd), are attributed to an aqueous fluid component. In some arcs and in Elliott’s worldwide compilation of ICP-MS data on arc basalts, this component has isotopic characteristics similar to hydrothermally altered MORB (e.g., 87Sr/86Sr , 0.7035, 143 Nd/144Nd , 0.5132; 208Pb/204Pb down of 38). High boron concentrations, and high boron and oxygen isotope ratios in the aqueous fluid component, relative to the MORB source, are attributed to hydrothermal alteration in the fluid source (Bebout et al., 1993, 1999; Domanik et al., 1993; Eiler et al., 2000; Ishikawa and Nakamura, 1994; Leeman, 1987, 1996; Ryan and Langmuir, 1992; Ryan et al., 1989; You et al., 1993, 1995). Our data compilation includes few data on boron and concentration, or on Boron and oxygen isotopes, and so we have not attempted to re-visit these topics. In contrast, enrichments in “fluid immobile” Th relative to light REE (e.g., Th/La) and enrichments of light/middle REE (e.g., La/Sm) are attributed to a partial melt of subducted material. This component is thought to be a sediment melt because in some arcs it has isotopic characteristics similar to subducting sediment (e.g., 87 Sr/86Sr up to 0.706; 143Nd/144Nd down to 0.5127; 208Pb/204Pb , 39.0) and because thorium enrichments at a given 143Nd/144Nd are larger than

623

can be accounted for by simple mixing of mantle peridotite and sediment. The fact that basalt mainly contributes aqueous fluid whereas the sediment rich component is mainly transported in melt might be attributed to the steep thermal gradient at the top of a subducting plate, in which only the topmost layers are heated above their fluid-saturated solidus, giving rise to the aphorism “sediments melt, basalts dehydrate.” The role of subducted sediment is particularly well documented for selected, high quality data on arc basalts in which the high Th/La component has Th/La identical to that in the subducting sediment column (Plank, 2003). Very efficient recycling of subducted thorium, together with subducted 10Be (present only in surficial sediments), has also been taken as evidence for transport of sediment-derived thorium and beryllium in a partial melt, rather than an aqueous fluid (e.g., Johnson and Plank (1999), Kelemen et al. (1995a); but apparently in disagreement with Morris (see Chapter 2.11) and Schmidt and Poli (see Chapter 3.17). (ii) Two main components? Using traceelement abundance. In our data compilation, there are many primitive lavas whose traceelement abundance is minimally affected by crystal fractionation, so we decided to dispense with the use of ratios such as Ba/La and Th/La. We did this because many arc lavas are enriched in both light REEs (the denominator) and barium, thorium, lead, and strontium (the numerator) in commonly used trace-element ratios. Thus, for example, a lava with 1 ppm lanthanum, Ba/La of 100, and Th/La of 0.1 has less lanthanum, barium, and thorium than a lava with 50 ppm lanthanum, Ba/La of 20, and Th/La of 0.3. Thus, in Figures 12 and 13 we plot incompatible trace-element abundances in primitive lavas. These data present a somewhat different picture from the three component hypothesis outlined in Section 3.18.3.3.1. In Figure 13, one can see that oceanic lavas with high Ba/La generally have lower 87Sr/86Sr than lavas with high Th/La, and that in most oceanic arcs, lavas with high Ba/La have low Th/La and lavas with high Th/La have low Ba/La, as outlined in Section 3.18.3.3.1. However, almost all data for the oceanic Aleutian arc have low 87Sr/86Sr (,0.7035) despite a wide range of Ba/La and

Figure 13 Trace-element concentrations in ppm, trace-element ratios, and isotope ratios for primitive arc lavas in our compilation. A high Ba/La, low Th/La component has 87Sr/86Sr ,0.704, while a distinct low Ba/La, high Th/La component in most oceanic arcs has 87Sr/86Sr , 0.706. However, continental arcs, the oceanic Aleutian arc, and some other primitive arc lavas have relatively high Ba/La and Th/La. For most oceanic arcs, Ba and Th concentration in primitive lavas are correlated (Figure 12) and both are correlated with 87Sr/86Sr. With very few exceptions, primitive andesites from the oceanic Aleutian arc and from continental arcs show lower 87Sr/86Sr and 208Pb/204Pb, at a given Ba or Th concentration, compared to lavas from other oceanic arcs at the same Ba or Th concentration. Symbols as in Figures 1 and 6.

624

Geochemistry of Subduction-related Magmatic Arcs

Th/La, and the Aleutians show strong positive correlation between Ba/La and Th/La. Figure 13 also shows that most oceanic arc lavas show positively correlated barium, thorium, lanthanum, and 87Sr/86Sr, while oceanic Aleutian arc lavas, and continental arc lavas, show wide variation in barium, thorium, and lanthanum that is not correlated with 87Sr/86Sr or, e.g., 208Pb/204Pb. Thus, for example, enrichments in thorium can occur in primitive arc lavas without incorporation of radiogenic strontium and lead, (and unradiogenic neodymium) from subducted sediments (Kelemen et al., 2003b). High thorium in primitive arc magmas with MORB-like lead, neodymium, and strontium isotope ratios, such as in primitive andesites from the western Aleutians, is probably not attributable to incorporation of a sedimentary thorium component, yet it has very similar thorium enrichment (though lower Th/La) compared to magmas, which are thought to incorporate thorium derived from subducted sediment. However, experimental studies of eclogite/aqueous fluid and peridotite/aqueous fluid partitioning show that thorium is not very soluble (Ayers et al., 1997; Brenan et al., 1996, 1995a,b; Johnson and Plank, 1999; Kogiso et al., 1997; Stalder et al., 1998; Tatsumi and Kogiso, 1997). In contrast, thorium is strongly enriched in small degree melts of basalt or sediment (e.g., Johnson and Plank, 1999; Rapp et al., 1999; Ryerson and Watson, 1987). For this reason, we attribute thorium enrichment in arc magmas generally to incorporation of thorium in a partial melt of eclogite facies, subducted basalt and/or sediment (e.g., Kelemen et al., 1993, 2003b). We return to this topic in Section 3.18.3.3.1. Somewhat to our surprise, over the past few years several investigators have found that small degree partial melts of eclogite may also have some characteristics normally attributed to an aqueous fluid component in arc magmas, such as high Ba/La, Pb/Ce, and Sr/Nd (e.g., Kelemen et al., 2003b; Rapp et al., 1999; Tatsumi, 2000). Thus, it may be difficult to distinguish between partial melts of eclogite and an aqueous fluid component using these traceelement ratios. Although we do not doubt that—for instance—barium, lead, and strontium can be fluid mobile, fluid enrichment of the arc magma source may not be the dominant control on barium, lead, and strontium enrichment in arc magmas in our compilation. Correlation of barium, lead, and strontium with thorium suggests that either “fluid” and “melt” components generally combine in similar proportions world wide, or that some process can produce both barium and thorium enrichment simultaneously. This observation and its interpretation are very similar to observations and

interpretation of 10Be versus B/Be systematics (e.g., Morris et al., 1990). Most primitive arc lavas have elevated barium, thorium, lanthanum, lead, caesium, strontium, and neodymium, compared to MORBs. Because all of these components are incompatible during melting of eclogite or peridotite, while they are fractionated from each other by fluid/rock partitioning, it may be that the difference between arc lavas and MORBs is primarily due to incorporation of a partial melt of subducted, eclogites facies sediment and/or basalt into the arc magma source. Where sediments are present, high sediment thorium and lead concentrations, together with distinctive sediment isotope ratios, are likely to impart a sedimentary signature to melts of subducted material. Where sediments are absent, Th/La and Pb/Ce are likely to be somewhat lower, and isotope ratios in subduction zone melts are likely to be closer to those in the MORB source. We do not mean to imply that dehydration reactions in subducting material do not evolve aqueous fluid. However, much of that fluid may escape into the fore arc region. And, certainly, scatter in Figure 12—with logarithmic axes—is indicative of a wide range of Ba/Th, Pb/Ce, and Sr/Nd. These variations may result from additional enrichment of barium, lead, and strontium (and boron) relative to thorium and light REE in some arcs due to an aqueous fluid component, as described in Section 3.18.3.3.1. (iii) Melt and fluid compositions and arc mass balance. As already mentioned, it is particularly important to consider enrichments in lanthanum and thorium, since these are relatively insoluble in aqueous fluids under subduction zone conditions (Ayers et al., 1997; Brenan et al., 1996, 1995a,b; Johnson and Plank, 1999; Kogiso et al., 1997; Stalder et al., 1998; Tatsumi and Kogiso, 1997). Given likely values for arc magma flux, average arc lava concentrations of these elements, assumptions about arc lower crustal composition, experimentally measured solubilities for these elements in aqueous fluids, estimates of the subducting flux of H2O, and assumptions about the mantle source of arc magmas, it has often been concluded that aqueous fluids derived from subducting material cannot explain the magmatic flux of lanthanum and thorium in arcs (e.g., Elliott, 2003; Elliott et al., 1997; Hawkesworth et al., 1997; Johnson and Plank, 1999; Kelemen et al., 1993, 2003b; Plank, 2003; Plank and Langmuir, 1988, 1993, 1998). If true, this is crucial because it requires that almost all arcs require enrichment of the mantle source in lanthanum and thorium via some mechanism other than aqueous fluid transport from subducting material. Because this is so crucial, we offer our own flux calculations in Table 2. These calculations

Characteristics of Arc Magmas Table 2(a)

Arc inputs and outputs assuming 5 wt.% H2O or 5 wt.% melt extracted. Aleutians

Age Thickness arc crust Material subducted Magmatic flux Magmatic flux Excess La Excess Th Excess Ba Excess Sr Excess Pb La flux in fluid La flux in fluid Th flux in fluid Th flux in fluid Ba flux in fluid Ba flux in fluid Sr flux in fluid Sr flux in fluid Sr flux in fluid Pb flux in fluid Pb flux in fluid La flux in melt Th flux in melt Ba flux in melt Sr flux in melt Pb flux in melt

625

my km 106 kg km21 yr21 106 kg km21 yr21 km3 km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21

55 20 1,321 115 38 292 117 20,809 38,016 396 133 679 27 52 99,692 47,136 11,827 33,566 37,606 861 957 2,689 356 72,557 79,353 573

Marianas

Izu-Bonin

Kermadec

Tonga

ref. D

45 20 592 140 47 155 31 7,224 16,601 83 67 345 10 20 15,605 7,378 5,215 14,800 16,581 323 360 1,364 134 11,358 34,988 215

45 20 896 140 47 95 209 5,023 12,361 542 106 541 16 31 22,725 10,745 7,738 21,963 24,606 547 609 2,143 209 16,540 51,922 364

30 18 1,154 150 50 503 166 19,823 29,106 416 131 673 22 42 37,121 17,551 10,152 28,813 32,281 1,175 1,307 2,662 287 27,017 68,117 782

24 12 1,564 113 38 569 67 29,367 38,075 619 171 876 23 44 37,576 17,766 13,616 38,646 43,297 1,399 1,556 3,465 300 27,348 91,363 931

1 2 1 3 3 2 1 3 2 1 3 4 4 4 4 4

Arc ages from Jarrard (1986), except Aleutians which is from Scholl et al. (1987). Material subducted is calculated assuming 7 km thick subducting oceanic crust with desities of 3.0 g/cm3; sediment thicknesses and densities from Plank and Langmuir (1998); subduction rates from England et al., 2003. Magmatic flux is calculated using arc crust thickness from Holbrook et al. (1999) (Aleutians), Suyehiro et al. (1996) (Izu-Bonin) and Plank and Langmuir (1988) (Marianas, Kermadec and Tonga), subtracting 6 km preexcisting oceanic crust, and assuming a 150 km arc width. Excesses of selected trace elements are calculated subtracting abundances in N-MORB (Hofmann, 1988) from average abundances in primitive arc basalt (Table 1). Trace element fluxes in fluid and melt are calculated assuming modal batch melting (5% in Table 2(a), 2% in Table 2(b)). Because estimates of partition coefficients (Ds) between eclogite and fluid are highly variable, more than one value was used for each element. References for the values used are listed in the column labelled Ref. D. Partition coefficients are from 1: Ayers (1998), 2: Stalder et al. (1998), 3: Brenan et al. (1995b), 4: Kelemen et al. (2003b). Average composition for Izu Bonin primitive basalt is not included in Table 1 (too few analyses) and is (in ppm): La ¼ 4.57, Th ¼ 1.68, Ba ¼ 49.75, Sr ¼ 201.49, Pb ¼ 4.36).

Table 2(b)

La flux in fluid La flux in fluid Th flux in fluid Th flux in fluid Ba flux in fluid Ba flux in fluid Sr flux in fluid Sr flux in fluid Sr flux in fluid Pb flux in fluid Pb flux in fluid La flux in melt Th flux in melt Ba flux in melt Sr flux in melt Pb flux in melt

Arc inputs and outputs assuming 2 wt.% H2O or 2 wt.% melt extracted.

kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21 kg km21 yr21

Aleutians

Marianas

Izu-Bonin

Kermadec

Tonga

Ref. D

52 284 11 22 99,012 25.678 4,811 15,009 17,131 621 764 1,466 264 50,570 44,852 321

26 144 4 8 15,499 4,019 2,121 6,618 7,553 233 287 744 99 7,916 19,776 121

42 226 6 13 22,570 5,853 3,148 9,820 11,209 395 486 1,168 155 11,528 29,347 204

52 281 9 17 36,868 9,561 4,130 12,883 14,705 848 1,044 1,451 212 18,830 38,501 438

67 366 9 18 37,320 9,679 5,539 17,280 19,723 1,010 1,243 1,889 223 19,061 51,640 522

1 2 1 3 3 2 1 3 2 1 3 4 4 4 4 4

Data sources and calculations as in Table 2(a). Concentration ranges in fluids and concentrations in melts shown in Figure 12, are calculated as flux divided by water/melt mass.

support previous work on the subject. Using experimental eclogite/fluid distribution coefficients, 2 –5% fluid equilibrating with the entire section of subducted oceanic crust þ sediments

beneath an arc (e.g., see Chapter 3.17) could carry the entire excess magmatic flux (primitive arc basalt—primitive MORB) of barium, lead, and strontium. Some experimental data have

626

Geochemistry of Subduction-related Magmatic Arcs

lanthanum solubilities in subduction fluids just high enough to account for the excess magmatic lanthanum flux in arcs. However, our results are consistent with previous calculations showing that aqueous fluid transport cannot account for the excess magmatic thorium flux. Further, our simple calculations support the idea that aqueous fluid transport should result in large fractionations of barium from thorium, lead from caesium, and strontium from neodymium. Thus, calculated fluid compositions lie at high Ba/Th, Pb/Ce, and Sr/Nd compared to barium, thorium, lead, strontium, and light REE enriched primitive arc lavas (Figure 12). In contrast, transport in 2 –5% melt of basalt þ sediment in eclogite facies can account for excess magmatic flux of barium, thorium, lanthanum, lead, and strontium, and this mechanism will produce relatively small fractionations between these different elements. As a result, calculated melt compositions plot at the enriched end of the trend from MORB to barium, thorium, lead strontium, and light REE enriched primitive arc lavas (Figure 12). (iv) Fluids, melts, or goo above the solvus? The suggestion that aqueous fluids might play a minor role in arc magma genesis is apparently odds with interpretations of data on oxygen isotopes versus trace-element enrichment in several arcs (Eiler et al., 2000). These data were inferred to indicate that enrichment in arcs is via an aqueous fluid, with a relatively high O/Ti ratio, rather than a silicate liquid with a much lower O/Ti ratio. However, the Eiler et al. result applies mainly to their relatively large data set for the Vanuatu arc, and depends on the composition of depleted boninite magmas, and questionable assumptions about mantle source composition (e.g., initial Cr/(Cr þ Al) ¼ 0.1). Further, although they did make a successful model involving melt transport rather than fluid transport, Eiler et al. discounted the result because it was apparently at odds with thermal models that rule out melting of subducting material. In their successful melt transport model, the melt has 30 wt.% H2O, within the range of 25 –50 wt.% H2O inferred for aqueous fluid saturated melts at 3– 5 GPa (Dixon and Stolper, 1995; Dixon et al., 1995; Kawamoto and Holloway, 1997; Mysen and Wheeler, 2000). The suggestion that aqueous fluids may not play a key role in subduction zone petrogenesis may seem at first to be at odds with the decades-old inference that addition of H2O to the mantle wedge is one of the key causes of mantle melting and arc magmatism. However, this is certainly not what we wish to propose. Instead, as elegantly shown in calculations by Eiler et al. (2000) the effect of adding H2O to peridotite is very similar whether the added H2O is in an aqueous fluid or dissolved in silicate melt. Thus, assuming that the effect of other possible fluxes such as K2O is

second order, if 10% fluid-fluxed melting requires addition of ,1 wt.% fluid with ,90 wt.% H2O, then 10% melt-fluxed melting might require addition of ,3 wt.% melt with 30 wt.% H2O, or ,2 wt.% melt with 45 wt.% H2O. About 30 – 50 wt.% H2O in a silicate melt corresponds to molar H/Si ,3– 6, which raises the question, is a fluid-saturated melt at 3– 5 GPa more like an anhydrous melt, or more like an aqueous fluid? It is possible that the H2O-rich phase generated via dehydration reactions in subducting plates at 3– 5 GPa might form at conditions where there is no longer a solvus separating distinct melt and fluid phases (Bureau and Keppler, 1999; Keppler, 1996). In this interpretation, the differences in experimentally constrained partitioning behavior at subduction zone pressures, for example between fluid/eclogite (Ayers et al., 1997; Brenan et al., 1996, 1995a,b; Johnson and Plank, 1999; Kogiso et al., 1997; Stalder et al., 1998; Tatsumi and Kogiso, 1997) and melt/eclogite (Rapp et al., 1999), might arise as a result of the H2O/silicate ratio in a given experimental bulk composition, rather the existence of distinct melt and fluid phases. Alternatively, the traditional interpretation, in which distinct fluid and melt phases can be present in H2O-eclogite at, e.g., 3– 5 GPa and 700 – 900 8C, may well be correct. The presence or absence of a solvus between melt and fluid in equilibrium with eclogite at high pressure and moderate temperature is likely to be controversial for several years to come. Meanwhile, the message from our flux calculations (Table 2) and calculated fluid versus melt compositions (Figure 12) remains clear: It is easiest to understand the range of trace-element enrichment in arc lavas, relative to MORBs, if transport of barium, thorium, lead, strontium, and light REEs from subducted sediment and basalt is mainly in a phase whose partitioning characteristics are similar to those measured for relatively H2O-poor “melt”/rock, and different from those measured for relatively H2O-rich “fluid”/rock. 3.18.3.3.2

Tantalum and niobium depletion

Tantalum and niobium in arc magmas are depleted relative to REEs, so that Nb/La and Ta/La are lower than in the primitive mantle and in MORBs. Depletion of primitive arc magmas in tantalum and niobium relative to lanthanum and thorium (Figure 14) is ubiquitous (a few niobiumenriched lavas—e.g., Kepezhinskas et al. (1997)—form a distinct anomaly which will not be discussed in this chapter). For this reason, and because thorium, tantalum, and niobium are relatively immobile in low temperature alteration of basalts, low Nb/Th and Ta/Th have been used as discriminants between arc magmas and both

Characteristics of Arc Magmas

La/Nb

10.00

1.00

.01

.10

1.00

10.00

1.00

10.00

Th/Nb

Nb (ppm)

100.00

10.00

1.00 .01

.10 Ta (ppm)

Figure 14 Relative depletion of Nb relative to Th and La in primitive arc lavas compared to MORB. For samples without Nb analyses (mostly, Aleutian data), we estimated Nb ,17*Ta. Primitive andesites have among the highest Th/Nb and La/Nb in the compiled data. Nb, Ta, and La concentrations are all higher than MORB in some continental arc lavas; in most oceanic arc lavas, including Aleutian primitive andesites, Nb and Ta concentrations are similar to or less than in MORB, while La is enriched compared to MORB. Symbols as in Figures 1 and 6.

ocean island basalts and MORBs, to constrain the provenance of lavas where tectonic accretion has obscured their original setting (e.g., Pearce, 1982; Pearce and Peate, 1995). While tantalum and niobium are relatively depleted, compared to other incompatible elements, many primitive arc magmas have higher tantalum and niobium concentrations than MORB (Figure 14). This is particularly true of primitive, continental arc lavas, both basalts and andesites. Oceanic arc lavas tend to have tantalum and niobium concentrations as low as, or lower than MORBs, and elevated lanthanum concentrations compared to MORBs.

627

Depletion of tantalum and niobium relative to other incompatible elements in arc lavas has been ascribed to many processes (review in Kelemen et al. (1993)) including (i) crystal fractionation of Fe –Ti oxides in the crust, (ii) fractionation of titanium-rich, hydrous silicates such as phlogopite or hornblende in the mantle or crust, (iii) extensive, chromatographic interaction between migrating melt and depleted peridotite, (iv) the presence of phases such as rutile or sphene in the mantle wedge (Bodinier et al., 1996), (v) relative immobility of tantalum and niobium relative to REE and other elements in aqueous fluids derived from subducting material, (vi) inherited, low Ta/Th and Nb/Th from subducted sediment (Plank, 2003), and (vii) the presence of residual rutile during partial melting of subducted material. Primitive basalts have not been affected by extensive FeTi oxide fractionation in the crust, and they have magmatic temperatures too high for amphibole or biotite crystallization, ruling out (i) and (ii) in the previous paragraph. Chromatographic fractionation of niobium and tantalum from thorium and lanthanum requires melt/rock ratios ,1023, and thus—given estimated arc fluxes—requires that parental arc magmas react with the entire mass of the mantle wedge during ascent (Kelemen et al., 1993), so (iii) seems unlikely. Consistently low Nb/La and Nb/Th is observed in arc magmas, even in primitive basalts which are hot, and far from rutile saturation at mantle pressures (Kelemen et al., 2003b), eliminating (v) for most arc basalts. Finally, as noted above, lanthanum and thorium enrichment in arc magmas probably occurs via addition of a melt from subducting sediment or basalt in eclogite facies, and not via aqueous fluid metasomatism, so high La/Nb and Th/Ta does not arise as a result of fluid/rock fractionation (vi). In some arc lavas, it appears that niobium, tantalum, and lanthanum concentrations are equally enriched, compared to MORBs, and thus low Nb/La and Ta/La in these may be due to transport of all these elements in a melt of subducted, continental sediment without residual rutile (Johnson and Plank, 1999). If so, however, this raises the question of how low Nb/La and Ta/La originally formed in the continents. Suites without a strong signature of recycled sediment (many oceanic arc lavas, and particularly primitive Aleutian andesites) generally show lanthanum enrichment without niobium and tantalum enrichment, relative to MORBs. For these reasons, if a single explanation for niobium and tantalum depletion in primitive arc lavas is to be sought, we prefer hypothesis (vii), fractionation of niobium and tantalum from other highly incompatible elements via partial melting of subducting, eclogite facies basalt or

628

Geochemistry of Subduction-related Magmatic Arcs

sediment with residual rutile (Elliott et al., 1997; Kelemen et al., 1993; Ryerson and Watson, 1987; Turner et al., 1997). If this inference is correct, it follows that nearly all arc magmas include a component derived from partial melting subducted material in eclogite facies, with residual rutile. 3.18.3.3.3

U-series isotopes

There have been numerous recent papers and reviews on U/Th, U/Pa, and Ra/Th isotopic disequilibrium in arc lavas (Bourdon et al., 1999, 2000; Clark et al., 1998; George et al., 2003; Gill and Condomines, 1992; Gill and Williams, 1990; Newman et al., 1984, 1986; Reagan and Gill, 1989; Reagan et al., 1994; Regelous et al., 1997; Sigmarsson et al., 2002, 1990, 1998; Thomas et al., 2002; Turner et al., 2003, 2000a,b,c, 2001, 1997; Turner and Foden, 2001). This is a complicated topic, and we cannot provide sufficient background information to make it accessible to a nonspecialist. However, because these data have bearing on other topics covered in this chapter, we summarize our understanding of recent work in this section. An explanation of the basic principles governing U-series fractionation and isotopic evolution is given by Spiegelman and Elliott (see Chapter 3.14). Several recent papers have emphasized the presence of substantial 226Ra excess (over parent 230 Th) in arc lavas. In the Marianas and Tonga arcs, 226 Ra excess correlates with Ba/La, Ba/Th, and Sr/Th (George et al., 2003; Sigmarsson et al., 2002; Turner et al., 2003, 2000a,b,c, 2001; Turner and Foden, 2001). As a result, 226Ra excess is linked in these papers to transport of a fluid component from subducted material to arc volcanoes in less than a few thousand years. The argument that 226Ra excess is generated by deep, subduction zone processes is particularly compelling for lavas from southern Chile, in which 10 Be/9Be correlates with 226Ra excess, suggesting that a component derived from young, subducted sediment may reach arc volcanoes in one or two thousand years (Sigmarsson et al., 2002). In contrast, a data set on beryllium- and uranium-series isotopes from the Aleutian arc does not show correlation of 226Ra excess and 10Be/9Be (George et al., 2003). Unfortunately Sigmarsson et al. (2002), do not report sufficient geochemical data to evaluate whether Ba/La is high in lavas with high 226 Ra excess, as in other data sets, and/or whether Th/La and Th/Ba are high, as in the proposed sediment melt component (Section 3.18.3.3.1) which might transport 10Be. If the high 226Ra component in the southern Andes has high barium and thorium, we would propose that radium, barium, thorium, and beryllium are all carried in a partial melt of subducted sediment or basalt, not in an aqueous fluid.

Transport of fluid and/or melt from the subduction zone directly beneath an arc to the surface in approximately one half-life of 226Ra requires transport rates of order ,100 m yr21, and this cannot be sustained during diffuse porous flow of melt through peridotite at porosities less than ,0.03. Instead, melt flow must be focused into high porosity conduits or cracks (Sigmarsson et al., 2002; Turner et al., 2001). It has been claimed that velocities ,100 m yr21 require flow of melt in fractures rather than via porous flow (e.g., Sigmarsson et al., 2002). However, simple calculations show that if melt/fluid viscosity is ,2 Pa s, density contrast between mantle and fluid/melt is ,500 kg m23, and mantle grain size is between 4 mm and 10 mm, the velocity of buoyancy driven porous flow of melt through mantle peridotite will exceed 100 m/yr21 at porosities of 0.09 –0.035 (Kelemen et al., 1997a). Estimates of porous flow velocity depend on uncertain parameterizations of mantle permeability, poorly constrained grain size in the mantle wedge, and so on. Nonetheless, it is apparent that—even if all 226Ra excess arises from dehydration and/or partial melting in subduction zones at a depth of ,100 km, and 226Ra excess data require transport in less than one half life of 226 Ra (1,600 yr)—this result cannot be used to discriminate between transport in fractures versus focused flow in high porosity conduits. In addition, while the currently accepted interpretation may well be correct, the present understanding of U-series data in arcs is reminiscent of early work on 226Ra excess in MORB, in which it was suggested that 226Ra excess forms during the initial stages of decompression melting, ,100 km below the seafloor, and is transported to ridge lavas in less than a few thousand years (e.g., McKenzie, 1985; Richardson and McKenzie, 1994). Currently available data on young MORB show a negative correlation between 226Ra excess and 230Th excess (e.g., Sims et al., 2002). Because 230Th excess is probably formed by melting or melt/rock reaction involving garnet peridotite (e.g., McKenzie, 1985; Spiegelman and Elliot, 1992), garnet pyroxenite (e.g., Hirschmann and Stolper, 1996; Lundstrom et al., 1995), or fertile, high pressure clinopyroxene (e.g., Turner et al., 2000b; Wood et al., 1999), the negative correlation between 230Th and 226Ra excesses may be indicative of a role for shallow level processes in the generation of 226Ra excess (e.g., melt/rock reaction in the shallow mantle, interaction with lower crustal plagioclase (Jull et al., 2002; Lundstrom, 2000; Lundstrom et al., 2000, 1995, 1999; Saal et al., 2002; Spiegelman and Elliot, 1992; Van Orman et al., 2002). As a result, 226Ra excess may not be a reliable indicator of melt transport velocities in the mantle beneath mid-ocean ridges (see Chapter 3.14).

Characteristics of Arc Magmas Arc lavas, like MORB, also show a negative correlation between 226Ra excess and 230Th excess (or 1/238U excess), allowing for some 226Ra decay (George et al., 2003; Reagan et al., 1994; Sigmarsson et al., 2002; Turner et al., 2003, 2000a,c, 2001) (Figure 15). Since the generation of 230 Th excess probably involves garnet and/or fertile clinopyroxene at pressures of 2 GPa or more, we anticipate the evolution of theories in which 230Th excess forms deep, while shallow processes play a role in generating 238U and 226Ra excess in arc lavas as well as in MORB. However, because radium, thorium, and uranium concentrations are higher in most arc lavas than in MORB, shallow level processes capable of generating 238U and 226Ra excess in arcs may differ from those beneath mid-ocean ridges. Thus, reaction between ascending melt and anhydrous mantle peridotite may not be capable of generating large 226Ra and 238U excesses in primitive arc lavas (Thomas et al., 2002), and assimilation of young plagioclase may also be an unlikely explanation for 226Ra excess in arc lavas (e.g., George et al., 2003). However, steady state diffusive gradients involving radium- and barium-rich minerals such as phlogopite or biotite could produce 226Ra excess correlated with high Ba/La in melts Shallow processes? 7

MORB arcs

decay

3

226Ra

4

226Ra

(226Ra/230Th)

5

decay

6

2 1 0 0.4

0.6

0.8

1 238

(230Th/

1.2

Residual garnet? 1.4

U)

Figure 15 226Ra excess versus 230Th excess in MORB and arc magmas. MORB data compiled by Sims et al. (2002). Arc data compiled by Turner et al. (2003). 226Ra decay may explain very low 226Ra excess in some lavas. However, the highest 226Ra excess at any given (230Th/238U) shows a negative correlation with 230Th excess. In the MORB data, 230Th excess may arise at depths of 2 – 3 GPa, via U/Th exchange between melt and garnet, or melt and high pressure, fertile clinopyroxene compositions. Lavas with high 226Ra excess do not show this high-pressure signature, and so many hypotheses call upon shallow melt/mantle interaction, or even lower crustal processes, to explain the genesis of 226 Ra excess. Similar theories have been advanced to explain the arc data. See text for discussion and references.

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interacting with the upper mantle or lower arc crust (Feineman and DePaolo, 2002). In this interpretation, relatively high Sr/Th in Marianas and Tonga lavas with high 226Ra excess (Turner et al., 2003) might be due to interaction with both biotite and plagioclase. Because radium and barium are geochemically similar to each other, and very different from thorium and lanthanum, there may be other processes that result in enrichment of radium and barium relative to thorium and lanthanum. From this perspective, not all lavas with high Ba/La and Ba/Th necessarily record selective enrichment of barium via aqueous fluid metasomatism in the mantle source. If 226Ra excess were always the result of relatively shallow processes, then correlation of high 10Be/9Be with high 226Ra excess in southern Chilean lavas would imply a shallow source for 10 Be enrichment as well. Some of the Chilean samples analyzed for 226Ra (Sigmarsson et al., 2002) are evolved (Stern and Kilian, 1996), and not enough data are presented to determine whether any of the other Chilean samples are primitive, or not. While excess 226Ra and 10Be may be transported from subduction zone depths in some arcs, it may also be worthwhile to reevaluate the extent to which some evolved magmas interact with meteoric water, or assimilate alteration products, that have high 226Ra excess and high 10Be/9Be. It is not yet clear whether high thorium, primitive lavas worldwide have high 230 Th excess, or not. 230Th excess has been observed in a few arc lavas, notably primitive andesites from Mt. Shasta in the southern Cascades (Newman et al., 1986), and high Mg# andesites from the Austral Andes in southernmost Chile (Sigmarsson et al., 2002). This is consistent with the hypothesis that primitive andesites contain a substantial proportion of partial melt from a source rich in residual garnet, such as subducting, eclogite facies sediment and/or basalt (Section 3.18.3.2.5). Substantial 230Th excess is also observed in some primitive basalts from Central America (Thomas et al., 2002). If primitive basalts do not include a substantial eclogite melt component, then the 230Th excess in Central American basalts might reflect melting or melt/rock reaction in the presence of garnet in the mantle wedge. (Another alternative is that 230Th excess arises as a result of melting, or melt/rock reaction, in the presence of fertile, high-pressure clinopyroxene (Turner et al., 2000b; Wood et al., 1999).) Two factors favor a potentially large role for garnet in the mantle wedge, compared to the melting region beneath mid-ocean ridges. First, relatively high H2O fugacity in the mantle beneath most arcs lowers mantle solidus temperatures at a given pressure. Because of the positive pressure/temperature

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Geochemistry of Subduction-related Magmatic Arcs

slope of reactions such as pyroxene þ spinel ¼ olivine þ garnet, lower solidus temperatures lowers the minimum depth at which garnet can be stable on the arc mantle solidus (e.g., Gaetani and Grove, 1998). Second, theoretical considerations suggest both porous and/or fracture transport of melt in the mantle wedge to the base of arc crust might be diagonal, beginning at depths of 150 km or more beneath the back arc region (Davies, 1999; Spiegelman and McKenzie, 1987). Thus the maximum pressure of mantle melting beneath an arc may not be constrained by the depth to the subduction zone directly beneath the arc. Because the maximum pressure of melting might be as large or larger beneath arcs, and the minimum pressure of garnet stability on the mantle solidus is probably lower beneath arcs, the interval in which garnet could react with arc melts is probably larger beneath arcs compared to ridges. We will return to this point in Section 3.18.3.3.4. More generally, arc lavas show 238U excess, usually interpreted to be indicative of high U/Th solubility in aqueous fluids derived from subducting material. 226Ra excess is correlated with 238U excess in most data sets on historically erupted lavas. Thus, if 226Ra excess is formed at great depth, it is likely (though not required; see Thomas et al. (2002)) that 238U excess is also formed at depth. Conversely, shallow processes that fractionate thorium from radium could also fractionate thorium from uranium. 3.18.3.3.4

Geodynamic considerations

Theory and observation were in conflict, and theory seemed to get the better of it. Fenner (1937, p. 166)

Since the 1980s, most geochemists and petrologists working on arcs have considered partial melting of subducted material to be relatively unusual, and absent beneath most modern arcs. This is based on three main lines of reasoning. First, very few arc lavas show strong heavy REE/middle REE depletion, but such depletion is predicted and observed in melts of eclogite or garnet amphibolite (e.g., Gill, 1974, 1978; Kelemen et al., 2003b; Rapp et al., 1999). In addition, partial melts of subducted material are likely to be granitic (e.g., Johnson and Plank, 1999; Nichols et al., 1994; Rapp et al., 1999; Rapp and Watson, 1995), and close to H2O saturation at 3 GPa or more, with ,25– 50 wt.% H2O (Dixon and Stolper, 1995; Dixon et al., 1995; Kawamoto and Holloway, 1997; Mysen and Wheeler, 2000). In these ways, no arc magma resembles a fluid-saturated melt of sediment or basalt in eclogite facies. Second, because many dehydration reactions in the subducting plate may be complete by 2 GPa, it

is not evident that free fluid will be available to facilitate fluid-saturated melting (e.g., Davies and Stevenson, 1992; Peacock et al., 1994; Rapp and Watson, 1995). Third, thermal models for arcs published between 1980 and 2002 uniformly indicate that the top of the subducting plate in “normal” subduction zones (convergence rate . 0.03 m yr21, subducting oceanic crust older than 20 Ma) does not reach temperatures above the fluid-saturated solidus for metabasalt or metasediment (see reviews in Kelemen et al. (2003a), Peacock (1996, 2003), and Peacock et al. (1994)). Recently, all three of these lines of reasoning have been challenged. First, modeling of reaction between heavy REE depleted melts of eclogite facies basalt or sediment and mantle peridotite shows that heavy REE abundances rise to levels in equilibrium with spinel peridotite at melt/rock ratios less than ,0.1, while the light REEs and other highly incompatible elements are almost unaffected (Kelemen et al., 1993, 2003b). Thus, primitive arc lavas with flat, middle to heavy REE patterns at low abundance, and light REE enrichment, could be formed by reaction between a partial melt of subducted material and the overlying mantle wedge. In addition, anatectic, H2O-rich melts of subducting sediment or basalt will both decompress and heat up as they rise into the overlying mantle. This “super-adiabatic” ascent may enhance melt/rock reaction, leading to a net increase in the liquid mass (Grove et al., 2001, 2003; Kelemen, 1986, 1990, 1995; Kelemen et al., 1993, 2003b). Thus, major elements in hybrid melts may be primarily derived from the mantle wedge, while highly incompatible trace elements may reflect the original, eclogite facies residue. (Note that the conclusion of the previous paragraph seems to be somewhat at odds with Section 3.18.3.3.3, in which it was proposed that 226 Ra and 238U excesses might arise via some shallow process. If this were so, then presumably other incompatible element concentrations would also be affected by this shallow process. However, based on data from (Newman et al., 1984; Sigmarsson et al., 2002), we anticipate that primitive andesites with a substantial eclogite melt component will have 230Th excess and little or no 226Ra excess). Second, it is now apparent that continuous dehydration reactions involving hydrous phases in metasediment and upper oceanic crust with higher pressure stability than glaucophane plus extensive solid solution, such as lawsonite, chloritoid, phengite, and zoisite, provide a small but nearly continuous source of fluid from shallow depths to depths exceeding ,250 km (review in Chapter 3. 17). In addition, hydrous phases such as serpentine and talc in the uppermost mantle of the subducting

Characteristics of Arc Magmas plate have an extensive stability field extending to high pressure, and will continue to dehydrate due to conductive heating of the cold interior of the subducting plate to depths up to ,200 km (review in Chapter 3.17). Finally, aqueous fluid may not be wetting in eclogite facies assemblages with abundant clinopyroxene. H2O-rich fluid in clinopyroxenite does not become interconnected until fluid fractions exceeding 7 vol.% (Watson and Lupulescu, 1993). Thus, some of the H2O evolved by dehydration reactions may remain within the metamorphic protolith until melting increases the permeability and permits H2O dissolved in the melt to escape the subducting slab by porous flow (Kelemen et al., 2003b), or until enough low density, interconnected aqueous fluid is present to fracture the overlying rock due to fluid overpressure (Davies, 1999). Thus, although some of the processes described in this paragraph are highly speculative, it is likely that through some combination of these processes, aqueous fluid is present to flux melting of subducting material at depths shallower than ,200 – 300 km. Third, computational and theoretical advances have made it possible for thermal models to incorporate temperature-dependent viscosity, and/or non-Newtonian viscosity, in the mantle wedge (Conder et al., 2002; Furukawa, 1993b; Furukawa and Tatsumi, 1999; Kelemen et al., 2003a; Kincaid and Sacks, 1997; Rowland and Davies, 1999; van Keken et al., 2002). This has several important effects, among them eliminating the necessity for prescribing the thickness of a rigid upper plate (Kelemen et al., 2003a). When the lithosphere in the upper plate is allowed to “find its own thickness,” this results in upwelling of the mantle to shallow depths near the wedge corner, so that asthenospheric potential temperatures extend to depths as shallow as 40 km beneath the arc. We think these models are preferable to previous, isoviscous models because they provide a much closer match to the high metamorphic temperatures recorded in exposures of arc Moho and lower crust (Kelemen et al., 2003a) and to calculated arc melt/mantle equilibration conditions (,1,300 8C at 1.2– 1.5 GPa, Bartels et al. (1991), Draper and Johnston (1992), Elkins Tanton et al. (2001), Sisson and Bronto (1998), Tatsumi et al. (1983)). They also provide an explanation for the observed, anomalously slow seismic structure observed in the uppermost mantle beneath arcs (e.g., Zhao et al., 1992, 1997), for the high heat flow in arcs (Blackwell et al., 1982; Furukawa, 1993b), and for the sharp gradient in the transition to very low heat flow in fore arcs (Kelemen et al., 2003a). The most recent of these models predict temperatures in the wedge and the top of the slab that are significantly higher compared to

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isoviscous models. Predicted temperatures are higher than the fluid-saturated solidus for both basalt and sediment (Johnson and Plank, 1999; Lambert and Wyllie, 1972; Nichols et al., 1994; Schmidt and Poli, 1998; Stern and Wyllie, 1973) near the top of the subducting plate directly beneath arcs at normal subduction rates and subducting plate ages (Kelemen et al., 2003a; van Keken et al., 2002). It would be premature to conclude that the tops of most subducting plates cross the fluid-saturated solidus directly beneath arcs, because this is an area of ongoing research. Also, it is not clear that subducting metasediment and/or metabasalt are fluid-saturated at these depths. And, even if fluids are present, natural rocks might evolve fluids with H2O activities lower, and melting temperatures higher, than in most melting experiments (Becker et al., 1999, 2000; Johnson and Plank, 1999). Nonetheless, it is clear that thermal models should no longer be invoked to “rule out” partial melting of subducted material in eclogite facies. Our community has been focused for a long time on very simple pictures of subduction. While this is expedient for maintaining sanity, it is intriguing to speculate briefly on the possibility of solid material transfer across the Benioff zone (Figure 16). Several decades ago, geochemists suggested that physical mixing of a few percent sediment with peridotite in the mantle wedge beneath arcs could account for many geochemical features of arc lavas (Armstrong, 1981; Kay, 1980). Indeed, by analogy with predicted gravitational instability of dense lower crust (e.g., Arndt and Goldstein, 1989; DeBari and Sleep, 1991; Ducea and Saleeby, 1996; Herzberg et al., 1983; R. W. Kay and S. M. Kay, 1988, 1991, 1993; S. M. Kay and R. W. Kay, 1985; Turcotte, 1989), subducting sediment may “delaminate” and rise into the overlying mantle wedge. Subducting sediment in eclogite facies is likely to be substantially less dense than the overlying mantle. Thus, viscous density instabilities will arise provided subducting sediment layers have thicknesses of 100 m to 1 km, once the overlying mantle viscosity becomes less than some critical value. For example, our calculations (Jull and Kelemen, 2001) suggest that, for density contrasts of 50– 150 kg m23 and background strain rates ,10214, a 1 km thick layer of subducting sediment would form unstable diapirs and rise into the overlying mantle at ,750 8C. This process will lead to mechanical mixing of sediment and mantle peridotite. In addition, rising, heating diapirs of sediment would certainly undergo partial melting in the mantle wedge. In another mechanism of solid transfer across the Benioff zone, imbrication of the subduction thrust at shallow depth, or downward migration of the subduction shear zone at greater depth, may transfer material from the top of the downgoing

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Geochemistry of Subduction-related Magmatic Arcs

Figure 16 Schematic illustration of mechanisms for transfer of sediments, volcanics, and/or lower crustal gabbros into the mantle wedge from the subducting plate and the base of arc crust. Dark black line indicates position of the “subduction zone”; below this line, material subducts at the convergence velocity. Above this line, material is carried downward more slowly. Any process leading to slow transport of low-melting point metasediment, metabasalt, or metagabbro into the mantle wedge would lead to partial melting of this material beneath an arc.

plate into the hanging wall, reducing its convergence velocity and allowing more time for conductive heating (Figure 16). “Subduction erosion” may transfer relatively hot, middle to lower crust from the fore-arc (metasediment, etc.) or the arc (garnet granulite, pyroxenite) to the top of the subducting plate (Clift and MacLeod, 1999; Ranero and von Huene, 2000; Vannucchi et al., 2001; von Huene and Scholl, 1991, 1993) (Figure 16). Because this material has a low melting point, compared to the mantle, but is already hot compared to the top of the subducting oceanic crustal section, conductive

heating will take it above its fluid-saturated solidus in a relatively short time, while fluid may supplied from the underlying, subducting plate. Similarly, although it is not “subduction,” delamination of dense crustal cumulates could lead to transformation of gabbroic rocks into eclogites, followed by heating, and partial melting (Gromet and Silver, 1987; R. W. Kay and S. M. Kay, 1993), as descending diapirs become entrained in ductile flow of the mantle wedge (Figure 16). Another important consideration is that partial melts of subducted material do not necessarily

Arc Lower Crust rise vertically through the mantle wedge from the subduction zone to the base of arc crust. One might expect initial trajectories to trend diagonally upward, away from the trench, due to the sum of buoyancy-induced upwelling and diagonal, downward solid flow in the mantle wedge. Higher in the wedge, theoretical considerations suggest that both porous flow and fracture transport of melt would be diagonally upward, toward the trench (Davies, 1999; Spiegelman and McKenzie, 1987). This diagonal, trenchward upwelling would be enhanced if solid flow in the upper part of the mantle wedge is also diagonally upward toward the trench, as is predicted from models incorporating a temperature dependent mantle viscosity (Conder et al., 2002; Furukawa, 1993b; Furukawa and Tatsumi, 1999; Kelemen et al., 2003a; Kincaid and Sacks, 1997; Rowland and Davies, 1999; van Keken et al., 2002). As a result, it may be a mistake to concentrate solely on whether subducted material exceeds its fluid-saturated solidus directly beneath an arc. Instead, if subducted material crosses the solidus at 150 km, or even 250 km, the resulting melt might be transported diagonally to the base of arc crust. In summary, for a variety of reasons, it is unwarranted to use geodynamical models to rule out partial melting of eclogite facies sediment, basalt, or lower crustal gabbro beneath arcs. Meanwhile, we believe that geochemical data on primitive arc lavas are best understood if a partial melt of subducted sediment and/or basalt (and/or gabbro) is the primary medium for transport of in compatible trace elements into the mantle wedge.

3.18.4 ARC LOWER CRUST Ever clearer became the danger of restricting attention to the ‘observed facts’. Direct observations are usually restricted to…a two-dimensional field. …Three dimensional it must be, in any case, even at the cost of one’s peace of mind—even at the cost of risking the quagmire of speculation about the invisible and intangible. There is, indeed, no other way. By declining Nature’s own invitation to think intensively about her third dimension, petrologists have ‘lost motion’ and have held back the healthy progress of their science. What petrology needs is controlled speculation about the depths of the Earth. Reginald A. Daly, Igneous Rocks and the Depths of the Earth (Daly, 1933) To reason without data is nothing but delusion. Arthur Holmes, The Age of the Earth (Holmes, 1937, p. 152)

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3.18.4.1 Talkeetna Arc Section Although there are outcrops of middle to lower crustal plutonic rocks with arc provenance in many places (e.g., DeBari, 1994; DeBari et al., 1999; Pickett and Saleeby, 1993, 1994), in this part of the chapter we will concentrate on data from the Talkeetna (south-central Alaska) and Kohistan (Pakistan Himalaya) arc crustal sections (e.g., Bard, 1983; Barker and Grantz, 1982; Burns, 1985; Coward et al., 1982; DeBari and Coleman, 1989; DeBari and Sleep, 1991; Jan, 1977; Miller and Christensen, 1994; Plafker et al., 1989; Tahirkheli, 1979). Both regions preserve tectonically dissected but relatively complete sections, from volcanics and sediments at the top to residual mantle peridotites near the base. Metamorphic equilibria at the Moho in both sections record conditions of ,1,000 8C and 1 –1.2 GPa. In addition, high P/T metamorphism, with pressure perhaps as high as 1.8 GPa, may be recorded by the Jijal complex, along the Main Mantle thrust at the base of the Kohistan section (Anczkiewicz and Vance, 2000; Ringuette et al., 1999; Yamamoto, 1993) and by late veins in the Kamila amphibolites (Jan and Karim, 1995); this is probably related to continental collision and exhumation of high P rocks along the Indian – Asian suture zone (Gough et al., 2001; Treloar, 1995; Treloar et al., 2001). Because we are most familiar with the Talkeetna section, we will emphasize data from our recent studies there, with supporting data from the Kohistan section. The Talkeetna section represents an arc fragment, ranging in age from ,200 Ma to ,175 Ma that was accreted along the North American margin and is now exposed in south central Alaska and along the Alaska Peninsula (Barker and Grantz, 1982; Detterman and Hartsock, 1966; Grantz et al., 1963; Martin et al., 1915; Millholland et al., 1987; Newberry et al., 1986; Nokleberg et al., 1994; Palfy et al., 1999; Plafker et al., 1989; Rioux et al., 2002b, 2001b; Roeske et al., 1989). The general geology and petrology of the Talkeetna section has been summarized by (Burns, 1983, 1985; Burns et al., 1991; DeBari, 1990; DeBari and Coleman, 1989; DeBari and Sleep, 1991; Newberry et al., 1986; Nokleberg et al., 1994; Pavlis, 1983; Plafker et al., 1989; Winkler et al., 1981). It is bounded to the north, along a contact of uncertain nature, by the accreted Wrangellia terrane. To the south, the Talkeetna arc section is juxtaposed along the Border Ranges Fault with accretionary wedge me´langes, the Liberty Creek, McHugh and Valdez complexes. This major fault has been a thrust and a right lateral strike-slip fault. Although the Border Ranges Fault is near vertical at present, a flat-lying klippe of gabbroic rocks, almost certainly derived from the Talkeetna section, overlies the McHugh complex in the Chugach Mountains

634

Geochemistry of Subduction-related Magmatic Arcs

north of Valdez. This is called the Klanelneechina klippe. Preliminary data suggest no inheritance in 200 – 180 Ma zircons from Talkeetna plutonic rocks (Rioux et al., 2002a, 2001a). The small contrast between neodymium isotopes in Talkeetna gabbros with neodymium isotopes for Jurassic MORB resembles the small neodymium isotope difference between Marianas arc lavas and present-day MORB (Greene et al., 2003). Small bodies of metaquartzite and marble intruded by Talkeetna plutonic rocks contain little or no zircon, and are interpreted as pelagic sediments (J. Amato, personal communication, 2003), while amphibolite rafts have andesitic compositions and the traceelement signatures of arc magmas (e.g., high La/ Nb and Th/Nb; our unpublished data). Thus, recent work is consistent with the hypothesis that the entire Talkeetna arc section is composed of late Triassic to middle Jurassic rocks that formed in an intra-oceanic arc. After reconnaissance mapping and sampling (Barker et al., 1994; Barker and Grantz, 1982; Newberry et al., 1986), the volcanic section of the Talkeetna arc has received relatively little study until now. Our preliminary data agree with earlier estimates that the volcanics are 5 –7 km thick. High Mg# basalts are relatively common, though 11 of our 87 samples are high Mg# andesites. More evolved lavas range from mainly tholeiitic andesites through tholeiitic and calc-alkaline dacites, to calc-alkaline rhyodacites and rhyolites. The volcanics are underlain and intruded by felsic to gabbroic plutons. Excellent descriptions of the petrology and major-element composition of the Talkeetna lower crust (e.g., Burns, 1985; DeBari and Coleman, 1989; DeBari and Sleep, 1991) and our recent work (Greene et al., 2003) show that much of the section is composed of compositionally monotonous gabbronorites. Some, but not all, include abundant magnetite. Prismatic hornblende of obvious igneous origin is rare, though most samples have hornblende rims—probably of deuteric origin—around prismatic pyroxene crystals. Olivine is extremely rare, even at the base of the crustal section. Both Talkeetna and Kohistan lower crustal gabbroic rocks have very high aluminum contents, compared to gabbroic rocks from ophiolites and to the average composition of continental lower crustal xenoliths (Rudnick and Presper, 1990). Because they have such high aluminum, these rocks can form more than 30% garnet at garnet granulite facies conditions, and are thus denser than mantle peridotite at the same pressure and temperature (Jull and Kelemen, 2001). DeBari and Coleman (1989), building on previous mapping by Burns, Newberry, Plafker, Coleman and others, concentrated much of their

work in the Tonsina area, where several small mountains preserve a laterally continuous Moho section, with relatively mafic gabbronorites overlying a thin but regionally continuous horizon of mafic garnet granulite (orthogneiss), overlying ,500 m of pyroxenite (mostly clinopyroxenerich websterite), in turn overlying residual mantle harzburgite with ,10% dunite (Figure 17). The harzburgite, in turn, is bounded by the Border Ranges fault to the south. It is important to emphasize that gabbronorites, garnet granulites, pyroxenites, and harzburgites show “conformable,” high temperature contacts which extend across intermittent outcrop for several kilometers. In this area, garnet granulites record conditions of ,1,000 8C and 1 GPa (DeBari and Coleman, 1989; Kelemen et al., 2003a), indicating that at the time of garnet growth, the crustal thickness was probably ,30 km. Previous studies (Burns, 1985; DeBari and Coleman, 1989; DeBari and Sleep, 1991) and our recent work (Greene et al., 2003) have found that Talkeetna gabbronorites and lavas form a cogenetic, igneous differentiation sequence.

3.18.4.1.1

Geochemical data from the Talkeetna section

We turn now to new geochemical data on the Talkeetna section (averages in Table 3). As in pilot data from Barker et al. (1994), REE patterns in all but two lavas are relatively flat, just slightly more light REE enriched than MORB (Figure 18). Extended trace-element “spidergrams” show the distinctive characteristics of arc magmas, such as high Th/Nb, La/Nb, Ba/La, Pb/Ce, and Sr/Nd. Many lavas, particularly the ones with the highest abundance of incompatible trace elements, show low Eu/Sm, consistent with plagioclase fractionation, and low Ti/Dy, which is probably due, at least in part, to magnetite fractionation. Many of these lavas are pervasively altered. Nonetheless, the parallelism of most trace-element patterns suggests that few of the geochemical characteristics summarized here have been substantially modified by alteration. Felsic plutonic rocks, primarily calc-alkaline tonalites, have spidergram patterns similar to the lavas. The gabbroic rocks, mostly gabbronorites, have spidergram patterns that are remarkably similar to the lavas in many ways, with high Th/Nb, La/Nb, Ba/La, Pb/Ce, and Sr/Nd (Figure 19). While DeBari and Coleman (1989) emphasized the presence of a few gabbroic samples with REE patterns reminiscent of MORB, suggesting that they might represent older oceanic crust into which the Talkeetna arc was emplaced, the extended trace-element patterns of all our gabbroic samples show arc-like signatures. For example, Talkeetna

Arc Lower Crust 145˚10'

635

Garnet granulite Gabbronorite 145˚05' A'

1 km Volcanics

Pyroxenite

61˚35' Harzburgite

Border ranges fault

Tonalite, diorite, gabbro and mafic dikes Serpentinite melange

15 km

A

Pyroxenite Antigorite schist Harzburgite Garnet granulite Melange 1,500 Gabbronorite 1,000 A

Gabbronorite Pyroxenite intrusions

A'

Gabbronorite 30 km

Pyroxenite Harzburgite

Figure 17 Schematic cross-section of the accreted, Jurassic Talkeetna arc in south-central Alaska based on new data and compilations by Greene et al. (2003), together with a geologic map and cross-section of the Moho exposure at Bernard Mountain, in the Tonsina region of the Talkeetna arc section, simplified from DeBari and Coleman (1989).

gabbronorites have average Th/Nb of 0.4, substantially higher than in gabbroic rocks in the Oman ophiolite (averaging 0.1 in our unpublished data) and lower oceanic crust (averaging 0.05; Hart et al. (1999)). High Ti/Dy in many gabbronorite samples is indicative of the presence of cumulate magnetite, which fractionates titanium from dysprosium during igneous crystallization, and high Eu/Sm reflects plagioclase accumulation. Similarly, high Pb/Ce and Sr/Nd in the gabbroic rocks may arise in part from lead and strontium enrichments in parental melts, and in part from the presence of cumulate plagioclase. In order to constrain which additional characteristics of the Talkeetna gabbronorites are inherited from trace-element features in parental, arc magmas, and which characteristics reflect igneous fractionation, we divide the average Talkeetna gabbronorite composition by the average lava composition to yield a highly approximate bulk crystal/liquid distribution coefficient pattern (Figure 20). This shows that, despite complementary lead and strontium enrichments in the lavas, high Pb/Ce and Sr/Nd in the gabbroic rocks also reflect the presence of cumulate plagioclase. In addition, note that thorium and uranium are more incompatible than rubidium, barium and potassium in plagioclase. Spidergrams for mafic garnet granulites from the Tonsina area show no evidence for the presence of

cumulate igneous garnet, which would be reflected in high heavy REE/middle REE ratios (Figure 21). Instead, they are very similar to typical Talkeetna gabbronorites. Thus, following DeBari and Coleman (1989), we believe that the garnets in this area are entirely of metamorphic origin. In contrast, gabbronorite, two-pyroxene quartz diorite and tonalite in garnet granulite facies from the Klanelneechina klippe, which record ,700 8C and 0.7 GPa, do show evidence for the presence of cumulate, igneous garnet (Figure 22). These rocks all have evolved compositions (whole-rock Mg# less than 50). They could be cumulates from evolved melts crystallizing in the lower crust, and/or restites produced via partial melting of high temperature lower crust. Trace-element contents of pyroxenites from the Tonsina area are shown in Figure 23. Although data on the most depleted of these may be very imprecise, they show high Pb/Ce and Sr/Nd, and generally low niobium and tantalum relative to potassium, lanthanum, and thorium. All of these characteristics are probably inherited from equilibrium with primitive arc magmas. Mu¨ntener et al. (2001) emphasized the strikingly low Al2O3 contents (1 – 6.5 wt.%, averaging 3.4 wt.% in our data) in many Talkeetna and Kohistan pyroxenites, despite their high pressures of crystallization. These low Al2O 3 contents were not produced in experiments on hydrous,

Table 3 New geochemical data on the Talkeetna section. Tonsina pyroxenite Normalized SiO2 TiO2 Al2O3 FeO* MnO MgO CaO Na2O K2O P2O5 Mg# XRF data Ni Cr Sc V Ba Rb Sr Zr Y Nb Ga Cu Zn Pb La Ce Th

N

oxides (wt.%) 49.95 17 0.07 17 3.42 17 8.31 17 0.17 17 27.95 17 9.88 17 0.22 17 0.02 17 0.01 17 86 17

(ppm) 516 3239 31.5 106 15.5 1.1 18.4 6.2 2.8 1.8 4.2 43.9 43.4 0.6 6.6 2.9 1.2

Std. error

0.296 0.003 0.126 0.229 0.002 0.343 0.355 0.006 0.001 0.000 0.2

17 15 17 105 17 1.0 17 4 17 0.5 17 0.1 17 1.5 17 0.2 17 0.1 17 0.0 17 0.1 17 2.9 17 1.8 17 0.1 17 0.4 17 0.4 17 0.1

Talkeetna gabbronorite

N

Std. error

47.86 0.66 19.00 9.94 0.18 7.78 12.53 1.82 0.16 0.08 58

95 95 95 95 95 95 95 95 95 95 95

0.047 0.004 0.030 0.035 0.001 0.027 0.025 0.009 0.002 0.001 0.1

31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31

1 4 0.5 5 3 0.1 2 0.5 0.2 0.0 0.1 2 0.7 0.1 0.2 0.2 0.0

42 167 39.5 250 75 2.0 303 17.4 9.8 1.9 15.7 88 55.8 1.6 5.6 8.3 1.3

Tonsina N garnet granulite 46.55 0.79 18.66 10.04 0.19 7.86 14.58 1.08 0.18 0.07 58 41 188 52 295 55 3.2 225 18 16 1.9 16.2 152 60 2.3 9.7 11.3 1.2

6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6

Std. error

1.234 0.122 0.153 0.722 0.014 0.486 0.903 0.081 0.067 0.019 2.8 3 21 2 30 13 1.4 9 3 2 0.2 0.6 27 8 0.5 2.2 2.5 0.2

Klanelneechina N garnet diorite and tonalite 51.60 0.89 18.92 10.30 0.21 4.89 10.41 2.51 0.13 0.14 45 11 55 38.7 242 36 0.7 277 45 17 3.6 17.9 34 85 0.4 5.3 14 1.0

7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7

Std. error

Intermediate to felsic plutons

N

Std. error

0.880 0.086 0.352 0.248 0.005 0.187 0.275 0.111 0.033 0.013 0.6

68.54 0.50 15.19 4.06 0.11 1.67 4.66 4.38 0.76 0.13 43

28 28 28 28 28 28 28 28 28 28 28

0.210 0.011 0.057 0.092 0.003 0.035 0.063 0.025 0.019 0.003 0.4

3 10 1.4 11 3 0.2 4 11 2 0.4 0.4 2 2 0.2 0.8 2 0.2

4.8 6.3 21.4 94 519 17.8 235 112 31 4.2 15.3 25 48 3.0 10.1 24.4 2.0

13 0.2 13 0.4 13 0.9 13 7 13 17 13 0.7 13 5 13 4 13 1 13 0.2 13 0.2 13 4 13 2 13 0.1 13 0.9 13 0.9 13 0.1

Lavas, tuffs and volcano-clastic 59.64 0.86 16.53 7.78 0.19 3.57 6.30 3.82 1.03 0.18 43 17.2 45.2 29.3 168.0 409 15.2 249 92.0 28.8 4.2 16.5 45.1 96.6 3.3 9.8 23.3 1.6

N

Std. error

114 114 114 114 114 114 114 114 114 114 114

0.069 0.003 0.017 0.026 0.001 0.017 0.033 0.011 0.007 0.001 0.1

86 86 86 86 86 86 86 86 86 86 86 86 86 86 86 86 86

0.3 0.7 0.1 1.2 5 0.2 2 0.6 0.1 0.0 0.0 0.5 0.9 0.0 0.1 0.1 0.0

ICP-MS data (ppm) La 0.192 Ce 0.417 Pr 0.059 Nd 0.328 Sm 0.149 Eu 0.063 Gd 0.244 Tb 0.049 Dy 0.362 Ho 0.081 Er 0.236 Tm 0.034 Yb 0.219 Lu 0.035 Ba 4.254 Th 0.024 Nb 0.096 Y 2.030 Hf 0.052 Ta 0.012 U 0.008 Pb 0.456 Rb 0.271 Cs 0.256 Sr 13.0 Sc 35.4 Zr 1.49

16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16 16

0.019 0.040 0.005 0.025 0.009 0.003 0.013 0.002 0.017 0.004 0.011 0.001 0.009 0.001 0.357 0.002 0.011 0.095 0.004 0.001 0.000 0.038 0.010 0.022 1.4 1.1 0.13

1.153 2.901 0.466 2.625 1.020 0.474 1.350 0.253 1.701 0.367 1.020 0.150 0.941 0.151 65 0.071 0.399 9.472 0.385 0.028 0.043 1.212 1.744 0.469 300 43.7 10.7

31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31 31

0.040 1.565 0.095 3.734 0.014 0.605 0.076 3.594 0.027 1.519 0.008 0.654 0.033 2.160 0.006 0.410 0.041 2.793 0.009 0.616 0.025 1.716 0.004 0.243 0.024 1.500 0.004 0.232 3 43 0.004 0.152 0.019 0.749 0.240 15.728 0.014 0.463 0.001 0.050 0.003 0.064 0.031 1.525 0.131 3.617 0.017 0.607 2 229 0.5 57 0.5 13

6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6

0.377 0.769 0.103 0.552 0.209 0.070 0.282 0.052 0.356 0.080 0.232 0.033 0.212 0.033 13 0.043 0.190 2.143 0.100 0.012 0.022 0.373 1.472 0.166 10 2 4

2.985 7.616 1.161 6.133 2.140 0.941 2.667 0.484 3.167 0.673 1.861 0.273 1.714 0.271 27 0.097 1.882 17.415 0.978 0.116 0.048 0.649 0.633 0.202 285 45 37

7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7

0.611 1.475 0.197 0.911 0.258 0.077 0.274 0.046 0.281 0.055 0.146 0.021 0.124 0.019 4 0.037 0.468 1.450 0.247 0.036 0.016 0.038 0.199 0.025 4 1 11

7.815 17.358 2.387 11.506 3.650 0.966 4.264 0.777 5.092 1.110 3.174 0.481 3.104 0.510 516 1.716 2.407 30.467 3.381 0.174 0.877 3.331 16.850 1.301 228 20.4 108

13 0.331 13 0.714 13 0.096 13 0.450 13 0.148 13 0.024 13 0.174 13 0.033 13 0.220 13 0.048 13 0.138 13 0.021 13 0.134 13 0.021 13 18 13 0.101 13 0.131 13 1.356 13 0.143 13 0.009 13 0.044 13 0.085 13 0.702 13 0.064 13 4 13 1.0 13 4

7.490 16.825 2.360 11.573 3.702 1.095 4.260 0.776 5.083 1.097 3.094 0.462 2.951 0.473 434 1.268 2.407 29.107 2.610 0.147 0.669 4.260 13.732 0.932 248 26.9 84.3

42 0.103 42 0.219 42 0.028 42 0.127 42 0.037 42 0.007 42 0.041 42 0.008 42 0.050 42 0.011 42 0.032 42 0.005 42 0.032 42 0.005 42 10 42 0.023 42 0.040 42 0.303 42 0.038 42 0.003 42 0.012 42 0.083 42 0.313 42 0.046 42 4 42 0.3 42 1.3

638

Geochemistry of Subduction-related Magmatic Arcs Talkeetna volcanics 100

Concentration/MORB

10

1

0.1

0.01 Rb Th K Nb Ce Pr Nd Hf Eu Tb Dy Er Yb Y Ba U Ta La Pb Sr Zr Sm Gd Ti Ho Tm Lu Ni Talkeetna felsic plutonic rocks 100

Concentration/MORB

10

1

0.1

0.01 Rb Th K Nb Ce Pr Nd Hf Eu Tb Dy Er Yb Y Ba U Ta La Pb Sr Zr Sm Gd Ti Ho Tm Lu Ni

Figure 18 Extended trace-element diagrams (hereafter, spidergrams) for volcanics and felsic plutonic rocks from the Talkeetna arc section, south central Alaska. Concentrations are normalized to N-MORB (Hofmann, 1988). Bold red lines are average values from Table 3. Talkeetna lavas, and plutonic rocks interpreted as liquid compositions, are only slightly enriched in light REE compared to MORB, but show depletion of Nb and Ta, and enrichment of Pb and Sr, typical for arc lavas worldwide. Their trace-element contents are similar to, for example, lavas from the modern Tonga arc. Data from Greene et al. (2003) and our unpublished research.

fluid-undersaturated arc basalt at 1.2 GPa, but were reproduced in experiments on primitive arc andesite at the same conditions. However, preliminary trace-element and neodymium isotope

data on Talkeetna pyroxenites suggest that they equilibrated with parental melts like the tholeiitic basalts that formed the overlying gabbroic and volcanic sections of the arc.

Arc Lower Crust

639

Talkeetna gabbronorites 100

Concentration/MORB

10

1

0.1

0.01

0.001 Rb Th K Nb Ce Pr Nd Hf Eu Tb Dy Er Yb Y Ba U Ta La Pb Sr Zr Sm Gd Ti Ho Tm Lu Ni

Figure 19 MORB-normalized spidergrams for lower crustal gabbronorites in the Talkeetna arc section, southcentral Alaska. Bold red line is average from Table 3. In these cumulate gabbros, some of the Nb and Ta depletion, and Pb and Sr enrichment, is inherited from parental, arc magmas, but the pattern is modified by high Ti (and probably Nb and Ta) in cumulate magnetite and high Pb and Sr in cumulate plagioclase. Data from Greene et al. (2003) and our unpublished research.

3.18.4.1.2

Composition, fractionation, and primary melts in the Talkeetna section

DeBari and Sleep (1991) estimated the bulk composition of the Talkeetna arc crust in the Tonsina area by adding compositions of different samples in proportions determined by their abundance in outcrop. They then added olivine, to calculate a primary magma composition for the Talkeetna arc in equilibrium with mantle olivine with Mg# of 90. This approach requires several assumptions. For example, the structural thickness of the Tonsina section cannot be more than ,15 km, whereas the garnet granulites at the base of the section record an original thickness of ,30 km. Thus, DeBari and Sleep assumed that tectonic thinning was homogeneously distributed over the entire section. Also, DeBari and Sleep assumed that outcrop exposures provided a representative estimate of the proportions of different rock types. This could be questioned. Up-section from the Moho level exposures, outcrop exposures in the Tonsina area are poor due to subdued topography and the presence of a major, Tertiary sedimentary basin to the NE. However, our new data on the composition of the most primitive Talkeetna lavas, high alumina basalts with Mg# , 60, are very similar to the bulk composition of the crust derived by DeBari and

Sleep on the basis of observed rock compositions and proportions. In addition, DeBari and Sleep assumed that the entire crust could be derived from a single, parental magma composition. Many arcs have isotopically and compositionally heterogeneous primitive magmas. However, our preliminary analyses show remarkable homogeneity in initial 143Nd/144Nd ratios (Greene et al., 2003). Also, with two exceptions, the lavas show nearly parallel spidergram patterns (Figure 18). In all these respects, our new data are consistent with the first-order assumption that most of the Talkeetna arc section was derived from a single type of primitive magma. Thus, the bold first-order approach of DeBari and Sleep (1991) has been largely vindicated by more extensive data. In order to better constrain the possible proportions of different igneous rock types prior to tectonic thinning, we recently completed least-squares modeling of crystal fractionation, using observed compositions and proportions of minerals in Talkeetna gabbronorites to reproduce the liquid line of descent from the most primitive Talkeetna high alumina basalts to average Talkeetna andesites and basaltic andesites (Greene et al., 2003). A striking result is that this modeling requires extensive crystallization of high Mg# clinopyroxene which is, in fact, very rare in the Talkeetna gabbronorites. Modeling predicts that ,20 –30%

640

Geochemistry of Subduction-related Magmatic Arcs Talkeetna gabbronorite/lava 100 Talkeetna felsic pluton/lava Talkeetna gbn/Oman lower gb Oman lower gabbro/MORB

Concentration ratios

10

1

0.1

0.01 Rb Th K Nb Ce Pr Nd Hf Eu Tb Dy Er Yb Y Ba U Ta La Pb Sr Zr Sm Gd Ti Ho Tm Lu Ni

Figure 20 Spidergrams showing ratios of trace-element concentration in Talkeetna arc gabbronorites/lavas (approximating bulk crystal/liquid distribution coefficients), Talkeetna felsic plutons/lavas, Talkeetna gabbronorite (gbn)/lower crustal gabbros from the Oman ophiolite, and Oman lower crustal gabbros/MORB. The gabbronorite/lava “distribution coefficients” show the effects of cumulate plagioclase (high Pb/Ce, Sr/Nd, Eu/Gd, and high Rb, Ba and K relative to Th and U) and magnetite (high Ti/Dy and very slightly high Ta and Nb/La). Talkeetna intermediate to felsic plutons have trace-element contents virtually identical to those in lavas, supporting the idea that the plutons represent melt compositions. Talkeetna gabbronorites are richer in REE and other incompatible elements, except Pb, than Oman lower crustal gabbros. Higher Pb in Oman versus Talkeetna lower crust probably reflects the influence of sulfide/sulfate equilibria during igneous crystallization. Oman cumulate gabbros show very depleted, smooth REE patterns extending to low Nb, Ta, U, and Th. This depleted pattern is interrupted by marked enrichments in Eu, Sr, Pb, K, Ba, and Rb, all of which reflect the presence of cumulate plagioclase. Talkeetna data from Greene et al. (2003) and our unpublished research. Oman data are samples from the Khafifah crustal section, Wadi Tayin massif, analyzed by ICP-MS at the Universite´ de Montpellier by Marguerite Godard, from Garrido et al. (2001) and our unpublished research.

of the gabbroic lower crust should have clinopyroxene Mg#. 85. In contrast, Figure 24 illustrates that none of our gabbroic samples have such high clinopyroxene Mg#s. In addition, most primitive Talkeetna lavas have Mg#’s that are too low for Fe/Mg exchange equilibrium with mantle olivine and pyroxene (incorporating reasonable assumptions about oxygen fugacity and Fe2þ/Fe3þ in arcs). Following Mu¨ntener et al. (2001), we infer that many hydrous, primary arc magmas are in a reaction relationship with olivine, forming by reactions such as orthopyroxene þ clinopyroxene þ spinel ¼ olivine þ melt. Such melts are in equilibrium with olivine, but will not crystallize olivine upon isobaric cooling, and instead will crystallize websterites (two pyroxene pyroxenites). This is consistent with the observation that cumulate dunites (olivine Mg# , 90) are absent and olivine pyroxenites are rare in the Talkeetna section. Thus, to constrain the composition of a primary magma in equilibrium with mantle peridotite, we

performed pyroxenite addition calculations, using the observed phase proportions in Talkeetna pyroxenites. Approximately 20 – 30% crystallization of pyroxenites from a primary melt in Fe/Mg equilibrium with mantle peridotite was required to produce the most primitive Talkeetna basalts (Greene et al., 2003). In other words, this modeling predicts that ,20 – 30% of the arc section should be composed of pyroxenite.

3.18.4.2 Missing Primitive Cumulates: Due to Delamination The great volume of andesite... in the orogenic belts is often taken up as a serious objection against the idea of its derivation from basalt magma by fractionation. Kuno (1968, p. 165).

The large proportions of pyroxenite and primitive gabbronorite predicted by fractionation modeling of primitive arc magmas contrast

Arc Lower Crust

641

Tonsina garnet granulite, Talkeetna arc 100

Concentration/MORB

10

1

0.1

0.01

0.001 Rb Th K Nb Ce Pr Nd Hf Eu Tb Dy Er Yb Y Ba U Ta La Pb Sr Zr Sm Gd Ti Ho Tm Lu Ni

Figure 21 MORB normalized spidergrams for mafic garnet granulites from the Tonsina area, at the base of the Talkeetna arc section at ,1 GPa, 1,000 8C (DeBari and Coleman, 1989; Kelemen et al., 2003a). Bold red line is average from Table 3. These orthogneisses formed via metamorphic recrystallization of protoliths with major and trace-element contents identical to gabbronorites from higher in the Talkeetna section. Garnet in these rocks is metamorphic, as previously proposed (DeBari and Coleman, 1989). Data from our unpublished research.

Klanelneechina klippe, Talkeetna arc garnet diorite and tonalite 100

Concentration/MORB

10

1

0.1

0.01 Rb Th K Nb Ce Pr Nd Hf Eu Tb Dy Er Yb Y Ba U Ta La Pb Sr Zr Sm Gd Ti Ho Tm Lu Ni

Figure 22 MORB normalized spidergrams for garnet diorites and tonalites from the Klanelneechina klippe, recording lower crustal depths (, 0.7 GPa, 700 8C; Kelemen et al., 2003a) from the Talkeetna arc section. Bold red line is average from Table 3. These evolved rocks (Mg# , 50) all include cumulate, igneous garnet, as indicated by their high heavy REE contents. They probably record partial melting of older arc lithologies under lower crustal conditions. Data from our unpublished research.

642

Geochemistry of Subduction-related Magmatic Arcs Talkeetna pyroxenite 10

Concentration/MORB

1

0.1

0.01

0.001 Rb Th K Nb Ce Pr Nd Hf Eu Tb Dy Er Yb Y Ba U Ta La Pb Sr Zr Sm Gd Ti Ho Tm Lu Ni

Figure 23 MORB normalized spidergrams for pyroxenites from the Tonsina area, at the base of the Talkeetna arc section. Bold red line is average from Table 3. Trace-element concentrations in these rocks are low, and data for, for example, U and Th may be very imprecise. Nonetheless, it is apparent that the pyroxenites inherited high Pb/Ce and Sr/Nd, and low Nb and Ta relative to K, Th, and U, from parental arc magmas. Data from our unpublished research.

dramatically with the observed proportion of primitive gabbronorites and pyroxenites in the Tonsina area. We have found no gabbronorite or garnet granulite (and only a very small outcrop of plagioclase pyroxenite) with clinopyroxene Mg#s . 85. Moreover, the thickness of the pyroxenite layer, between overlying gabbroic rocks and underlying residual mantle peridotites, is less than ,500 m. The relationship between clinopyroxene Mg# and whole-rock Mg# in Talkeetna gabbronorites (Figure 24), together with our larger data set on whole-rock compositions, confirms the observations based on clinopyroxene analyses. Modeling predicts that there should be a large proportion of gabbronorites and pyroxenites with whole-rock Mg# . 80, but in fact there are very few. This observation is strikingly similar to compiled data on whole-rock Mg# in the Kohistan arc section, in which there are no gabbroic rocks with Mg# . 80, and pyroxenites with Mg# . 80 are mainly found in a narrow band, less than 3 km thick, immediately above the Moho. There are small, ultramafic intrusions within the gabbroic lower crust in both the Talkeetna and Kohistan arc sections. However, they comprise less than 5% of the outcrop area. Also, at least in the Talkeetna crustal section, these bodies generally have clinopyroxene Mg# , 85.

Thus, the proportions of igneous rocks calculated from modeling of the liquid line of descent in the lavas are strikingly different from those observed in the Talkeetna and Kohistan section. The bulk of the predicted, primitive cumulates, with clinopyroxene Mg# between 92 and 85, are apparently missing. This result is both uncertain and important, and so we provide additional constraints on the modeling here. The crystal fractionation modeling of Greene et al. (2003) requires assumptions about the Fe2þ/Fe3þ ratio, and about pyroxene/melt Fe/Mg equilibria, which are imprecise. For this reason, in Figure 25 we present alternative methods for estimating the proportion of primitive cumulates that were produced by crystal fractionation in the Talkeetna arc section. The left-hand panel of Figure 25 shows clinopyroxene Mg# in equilibrium with Talkeetna lavas versus ytterbium concentration in the same lavas. If ytterbium were a completely incompatible element, quantitatively retained in melts during crystal fractionation, then a doubling of the ytterbium concentration would indicate 50% crystallization. In fact, ytterbium is only moderately incompatible, so doubling of ytterbium indicates more than 50% crystallization. It can be seen from these data that a decrease in clinopyroxene Mg# from 85 to 75 is accompanied by more than 50% crystallization. If this trend can be

Arc Lower Crust

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Figure 24 Histograms of clinopyroxene Mg# in mafic plutonic rocks from the Talkeetna arc section, whole-rock Mg# in mafic plutonic rocks from the Talkeetna arc section, and whole-rock Mg# in mafic plutonic rocks from the Kohistan arc section, NW Pakistan. Also shown is average clinopyroxene Mg# versus whole-rock Mg# for samples from the Talkeetna section. Assuming that the number of samples analyzed is representative of the proportion of different compositions present in the Talkeetna and Kohistan sections, these data show that no gabbroic rocks in the Talkeetna section have clinopyroxene Mg# . 86 or whole-rock Mg# . 80, and only 3 of 100 gabbroic rocks in the Kohistan section have whole-rock Mg# . 80. Pyroxenites in the Tonsina area in the Talkeetna section, with clinopyroxene Mg# from 92 to 86 and whole-rock Mg# from 92 to 76, form a layer ,500 m thick between underlying residual peridotite and overlying garnet granulite and gabbronorite. Pyroxenites in the Kohistan section, with whole-rock Mg# from 92 to 72, mainly occur in a band a few km thick between underlying residual mantle peridotites and overlying garnet granulite and gabbroic rocks in the Kohistan section. More evolved pyroxenites form volumetrically insignificant intrusions into gabbroic rocks in both sections. Crustal thickness in both sections is estimated to have been ,30 km, based on metamorphic equilibria (see text). Thus, pyroxenites and primitive gabbroic rocks with Mg# from 92 to 85 comprise a very small proportion of both the Talkeetna and Kohistan arc sections. These observed proportions are in striking contrast to estimates based on fractionation models and experimental data requiring , 30% crystallization of primitive cumulates with clinopyroxene Mg# between 92 and 85. Talkeetna data from Greene et al. (2003) and our unpublished research. Kohistan whole-rock compositions from George et al. (1993), Hanson (1989), Jan (1988), Jan and Howie (1981), Jan et al. (1982), Khan et al. (1993, 1989), Miller and Christensen (1994), Petterson et al. (1991, 1993), Petterson and Windley (1985, 1992), Shah and Shervais (1999), Sullivan et al. (1993), Treloar et al. (1996), Yamamoto and Yoshino (1998), and Yoshino and Satish-Kumar (2001).

Experimental data and models

Talkeetna lava data 7

110

6

100

~30% xl'n cpx Mg# 92 to 85 (except olivine model) ~16% xl'n

90 Melt (wt.%)

5 Yb (ppm)

Initial cpx Mg# calc from bulk composition

4 3

80 Greene et al., Talkeetna model r 2 = 0.96 70 olivine fractionation only r 2 = 0.96 60 Muentener et al., 2001, experiments

2 > 50% xl'n ∆cpx Mg# = 10

1

50

85–44 y = –310 + 461 Mg# r 2 = 0.93

40

0 55

60

65 70 75 80 Clinopyroxene molar Mg#

85

90

30 75

85–41c y = –842 – 2,115 Mg# + 1,430 Mg# r 2 = 0.99 80

85 90 Clinopyroxene molar Mg#

95

Figure 25 Estimates of melt fraction versus clinopyroxene Mg# from a variety of different methods. On the left, Talkeetna arc lava data are used to establish an empirical relationship between measured Yb concentration, and calculated clinopyroxene Mg# based on a clinopyroxene/melt Fe/Mg Kd of 0.23 (Sisson and Grove, 1993a). If Yb were a perfectly incompatible element, entirely retained in the melt, doubling of the Yb concentration would reflect 50% crystallization. Since Yb is not perfectly incompatible, doubling of Yb indicates more than 50% crystallization. Thus, a change in clinopyroxene Mg# from ,85 to ,75 is associated with more than 50% crystallization (Filled symbols used in linear fit; open symbols omitted from fit). On the right, experimental and model results are used to constrain similar relationships. Data on experimental crystallization of hydrous primitive arc basalt (85-44) and primitive arc andesite (85 – 41c) at 1.2 GPa (Mu¨ntener et al., 2001) yield two empirical relationships between melt fraction and clinopyroxene Mg#, in which a change in clinopyroxene Mg# from 92 to 85 required ,30% crystallization. Least squares modeling of the Talkeetna liquid line of descent, based on Talkeetna lava, pyroxenite, and gabbronorite compositions, provides a similar result (Greene et al., 2003). Olivine is rare in Talkeetna pyroxenite and gabbroic rocks. Olivine in dunites and residual mantle harzburgites underlying the pyroxenites in the Tonsina area has Mg#s , 90 (in equilibrium with clinopyroxene Mg# , 92), so it does not record substantial crystal fractionation. Nonetheless, the right-hand panel in this figure also shows an olivine-only fractionation calculation, using an estimated Talkeetna primary melt (Greene et al., 2003), in terms of equivalent clinopyroxene Mg# versus melt fraction. This model provides a minimum bound of ,15% crystallization to shift clinopyroxene Mg# from 92 to 85.

Arc Lower Crust extrapolated to crystallization of more primitive melts, it suggests that the decrease in clinopyroxene Mg# from mantle values (,92) to typical Talkeetna gabbronorite values (,85) was produced by ,35% crystallization. The right-hand panel in Figure 25 illustrates results of experimental crystallization of pyroxenite and primitive gabbronorite from hydrous arc basalt and andesite at 1.2 GPa (Mu¨ntener et al., 2001). Again, these data indicate that ,30% crystallization is required for clinopyroxene Mg# to decrease from ,92 to ,85, consistent with the modeling of (Greene et al., 2003). There is no evidence for extensive fractionation of olivine fractionation from primitive Talkeetna magmas. Dunites are present in the mantle section, but these are probably replacive melt conduits (see Kelemen, 1990; Kelemen et al., 1997a,b; and references therein). In any case, mantle dunites have olivine Mg# . 90, and thus it is apparent that they do not record extensive crystal fractionation. Nonetheless, in Figure 25 we also illustrate a model for olivine fractionation alone. This yields a lower bound of ,15 wt.% olivine crystallization required to decrease melt Mg# so that equilibrium clinopyroxene Mg# is decreased from 92 to 85. Finally, our results are similar to those of earlier least-squares fractionation models, which require 21% (Conrad and Kay, 1984) and 16 – 26% (Gust and Perfit, 1987) crystallization of ultramafic cumulates to produce high-aluminum basalt in island arcs. There are at least four possible explanations for the discrepancy between the proportion of high Mg# cumulates inferred from crystal fractionation modeling and the proportion of high Mg# cumulates observed in the Talkeetna and Kohistan sections. (i) The exposed section is not representative of the original arc crustal section. In the Talkeetna area, outcrop is discontinuous, due to numerous faults and subdued topography. It is evident that the present-day structural thickness of the section (,20 km from Moho to volcanics) cannot be as great as the thickness inferred on the basis of thermobarometry, ,30 km. However, the missing section is unlikely to be high Mg# pyroxenites near the Moho, because the Tonsina area gabbronorites and garnet granulites with clinopyroxene Mg# , 85 are in high temperature, conformable contact with high Mg# pyroxenites, which in turn are interfingered along high temperature contacts with residual mantle peridotites. Thus, in this section, the original thickness of cumulates with clinopyroxene Mg# between 85 and 92 as exposed in continuous outcrop, is only ,500 m. (ii) The missing Talkeetna arc high Mg# plutonic rocks crystallized beneath the “Moho” exposed in the Tonsina area. While arc mantle is rarely exposed, worldwide, it is apparent from

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dredging at mid-ocean ridges and from ophiolite studies that gabbroic and ultramafic intrusions can form lenses within the residual mantle beneath oceanic spreading ridges (e.g., Cannat, 1996). In the Tonsina area of the Talkeetna section, only a 1– 2 km of residual mantle is exposed. Thus, we cannot rule out the presence of high Mg# pyroxenites emplaced as plutons or sills within mantle peridotites at greater depth. (iii) Equilibrium crystallization of primitive magmas (as distinct from fractional crystallization) occurred until ,70 – 50% of the initial liquid remained. The remaining liquid was then efficiently extracted, leaving relatively pure gabbronorite cumulates with clinopyroxene Mg#s less than 85. This might happen if, for example, dense, rising melts ponded near the Moho, underwent partial crystallization, and then less dense, evolved melts were extracted to form the overlying crust. Calculated densities for Talkeetna average gabbronorite at pressures from 0.1 GPa to 0.8 GPa are ,3,000 kg m23 (Jull and Kelemen, 2001). Calculated densities for the primary and primitive melts for the Talkeetna section (Greene et al., 2003) are ,2,800 kg m23 on an anhydrous basis, and would be less if—as we surmise—they contained several weight percent H2O. Thus, primitive Talkeetna melts would have been buoyant with respect to the igneous crust and the Moho would not have been a zone of neutral buoyancy. Alternatively, melts may have ponded at depth beneath a permeability barrier, in a transition from porous flow to flow in melt induced fractures, as proposed for the base of the crust at mid-ocean ridges based on observations in the Oman ophiolite (Kelemen and Aharonov, 1998; Kelemen et al., 1997b; Korenaga and Kelemen, 1997) and geophysical data from the East Pacific Rise (Crawford and Webb, 2002; Crawford et al., 1999; Dunn and Toomey, 1997). However, the idea that primary Talkeetna magmas underwent 30 – 50% equilibrium crystallization to form gabbronorites with clinopyroxene Mg#s , 85 is not consistent with the presence of plagioclasefree pyroxenites with pyroxene Mg#s from 92 to 85, nor with experimentally determined phase equilibria for hydrous, primitive arc magmas that indicate a substantial interval for plagioclase-free, pyroxenite crystallization. (iv) Gravitational instability of dense cumulates overlying less dense upper mantle peridotites may have induced viscous “delamination” at the base of the crust. This is consistent with the observation that the Mg# “gap” occurs near the Moho, and the few rocks with clinopyroxene Mg#s from 90 to 85 are pyroxenites. As pointed out by Arndt and Goldstein (1989), and quantified by Jull and Kelemen (2001) and Mu¨ntener et al. (2001), pyroxenites are generally denser than upper mantle peridotites, because olivine and pyroxene densities

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are very similar at similar Mg#, but igneous pyroxenites have lower Mg# than residual peridotites. Furthermore, the density contrast between ultramafic cumulates and mantle peridotites is sufficient to drive viscous “delamination” of a pyroxenite layer 1 – 2 km thick in ,10 Ma at ,750 8C and ,105 yr at ,1,000 8C (stress dependent, non-Newtonian olivine rheology at a background strain rate of 10214 s21; see Jull and Kelemen (2001) for details and other estimates). Although all four hypotheses outlined above remain possible, for the reasons outlined in the previous paragraphs we favor the fourth hypothesis, in which pyroxenites foundered into the underlying mantle as a result of density instabilities. Even if some of pyroxenites were emplaced into residual mantle beneath the Moho, the high temperature of sub-arc mantle, and the high strain inferred from Talkeetna peridotite ductile deformation fabrics, combined with the short instability times calculated by Jull and Kelemen (2001), suggest that any significant thickness of pyroxenite near the base of the crust or within the mantle would have been viscously removed during the duration of arc magmatism. Garnet granulites with the composition of Talkeetna gabbronorites would also be denser than the underlying mantle (Jull and Kelemen, 2001). We suggest that it is no coincidence that the “Mg# gap” coincides with the garnet-in isograd at the base of the Talkeetna gabbroic section. Thus, along with the missing pyroxenites, we infer that garnet granulites may also have been removed from the base of the Talkeetna arc section via density instabilities. In this view, the narrow horizons of pyroxenite and garnet granulite along the Moho in the Talkeetna section are a small remnant of a much larger mass of primitive cumulates, most of which were removed by delamination.

3.18.4.3 Garnet Diorites and Tonalites: Igneous Garnet in the Lower Crust In the Klanelneechina klippe, which was thrust south of the Talkeetna arc section over younger, accretionary complexes, we found that most rocks are pyroxene quartz diorites to tonalites (bulk Mg# , 50, plagioclase anorthite content ,50 mol.%), and include some garnet-bearing assemblages. Metamorphic equilibria in garnet þ pyroxene bearing rocks record ,7 GPa, 700 8C (Kelemen et al., 2003a). While these rocks will be discussed elsewhere in detail, they are pertinent to the question of continental genesis, and so we illustrate their trace-element contents in Figure 22. While their LILE, HFSE, and light REE patterns vary widely, these samples consistently show heavy REE enrichment, and low Ti/Dy. These characteristics are probably indicative of

the presence of cumulate or residual, igneous garnet. Felsic melts extracted from these rocks were garnet-saturated at a depth of ,20 km in the Talkeetna arc crust. They were light REE enriched, heavy REE depleted, high SiO2, low Mg# melts that could have mixed with primitive melts to produce high Mg# andesite compositions.

3.18.5 IMPLICATIONS FOR CONTINENTAL GENESIS 3.18.5.1 Role of Lower Crustal Delamination in Continental Genesis The similarity of trace-element patterns in arc magmas and continental crust has led to the inference that most continental crust is derived from igneous, arc crust. We know of three potentially viable explanations for the generation of andesitic, continental crust with an Mg# of ,0.5 via processes involving arc crust: (i) Crystal fractionation from, or partial melting of, a primitive arc basalt composition at high fO2 and high fH2O, forming high Mg# andesite melt and corresponding, low SiO2 cumulates, followed by delamination of the resulting solid residue. (ii) Magma mixing or simply juxtaposition of primitive arc basalt and evolved, granitic rocks, together with delamination of the solid residue left after granite generation. (iii) Crystal fractionation from, or partial melting of, a primitive arc andesite composition, followed by delamination of the corresponding cumulates. In (i) – (iii), delamination would be unnecessary if primitive cumulates or residues of lower crustal melting were ultramafic, and remained present below the seismic Moho. However, the absence of pyroxenite layers ,10 km in thickness in the Talkeetna and Kohistan arc sections suggests that large proportions of ultramafic, igneous rocks are not present at the base of arc crust. Thus, all three explanations for the genesis of continental crust require delamination of garnet granulite, eclogite and/or pyroxenite. Moreover, recent, dynamical calculations support the hypothesis that delamination is possible—even likely—where Moho temperatures exceed ,750 8C, where crustal thicknesses reach 30 km or more, and where ultramafic cumulates are present (Jull and Kelemen, 2001). The base of arc crust fulfills all of these criteria. Finally, our data on the Talkeetna arc section, and more limited data on the Kohistan arc section, support the hypothesis that substantial proportions of pyroxenite, and perhaps also garnet granulite, were removed by viscous delamination from the base of the arc crust.

Implications for Continental Genesis If delamination of dense lower crustal rocks has been essential to continental genesis, and the delaminated rocks constitute ,20– 40% of the mass of the continents, one might expect to see evidence for this component in magmas derived from the convecting mantle. However, continental crust comprises only ,0.5% of the silicate Earth, so that—even if it represents 40% of the original crustal mass—recycled lower crust might comprise a very small fraction of the convecting mantle. Tatsumi (2000) argued that recycled pyroxenite or cumulate gabbro, added to the convecting mantle via delamination from the base of arc crust, has produced the EM I isotopic end-member observed in some ocean island basalts, because he inferred on the basis of modeling that this recycled material would have high Rb/Sr, U/Pb, and Th/Pb, and low Sm/Nd compared to bulk Earth and the primitive mantle. However, although the Tonsina pyroxenites have variable trace-element patterns, the average and median compositions have U/Pb and Th/Pb ratios less than in MORBs and primitive mantle, and average Sm/Nd and Lu/Hf greater than in MORBs and primitive mantle. In general, one might expect delaminated Talkeetna pyroxenite to evolve isotope ratios similar to, or more depleted than, the MORB source. Such compositions could, in fact, be present in the source of the more depleted end-members of the MORB isotope spectrum. If a delaminated lower crustal component included substantial amounts of garnet granulite with the composition of garnet granulites or gabbronorites from the Tonsina region in the Talkeetna arc section, the trace-element ratios and abundances in this reservoir would be different from those in delaminated pyroxenite. However, it remains true that Rb/Sr, U/Pb, and Th/Pb in most of our garnet granulite and gabbronorite samples are lower than in the MORB source, while Sm/Nd and Lu/Hf are higher than in the MORB source. Again, as a consequence, long term isotopic evolution followed by melting of this component during upwelling in the convecting mantle, would yield a melt with isotope ratios similar to, or more depleted than, the MORB source. While incompatible traceelement concentrations in pyroxenite are really very low, trace-element concentrations in the garnet granulites and gabbronorites are comparable to those in MORBs. Thus, given the large strontium and lead anomalies in the garnet granulites, removal of substantial amounts of garnet granulite from the base of continental crust would tend to decrease Pb/Ce and Sr/Nd in the remaining crust. This provides one possible explanation for the fact that Sr/Nd in continental crust is lower than in otherwise geochemically similar arc magmas, as also noted by Kemp and Hawkesworth (see Chapter 3.11) and Rudnick (1995).

647

3.18.5.2 Additional Processes are Required The process of delamination—if it occurred— apparently did not produce an andesitic bulk composition in the remaining Talkeetna arc crust. Instead, the Talkeetna arc section probably has a basaltic bulk composition (DeBari and Sleep, 1991; Greene et al., 2003) even after removal of dense, primitive cumulates. This inference is consistent with seismic data, and the composition of primitive arc basalts from the central Aleutian arc and the Izu – Bonin – Marianas arc system, all of which appear to have bulk crust compositions that are basaltic rather than andesitic (Fliedner and Klemperer, 1999; Holbrook et al., 1999; Kerr and Klemperer, 2002; Suyehiro et al., 1996). As noted above, some additional processes must be required to produce continental crust.

3.18.5.2.1

Andesitic arc crust at some times and places

Our favored hypothesis is that continental crust was mainly produced by fractionation of olivine and clinopyroxene from primitive andesite. Following many others (Defant and Kepezhinskas, 2001; Drummond and Defant, 1990; Martin, 1986, 1999; Rapp and Watson, 1995; Rapp et al., 1991), we believe that higher mantle temperatures in the Archean led to more common, larger degrees of partial melting of subducting, eclogite facies basalt. Alternatively, or in addition, dense lower crustal rocks foundering into the hot upper mantle may have commonly undergone partial melting (Zegers and van Keken, 2001). Also, due to higher degrees of melting at hot spots and beneath spreading ridges, more depleted peridotite was present in the upper mantle, including the mantle wedge above Archean subduction zones. Thus, interaction between eclogite melts and highly depleted mantle peridotite yielded primitive andesite (e.g., Kelemen et al., 1998; Rapp et al., 1999; Ringwood, 1974; Rudnick et al., 1994), but produced little additional, basaltic melt. Delamination removed mostly ultramafic cumulates from the base of the crust. Intracrustal differentiation of an andesitic bulk composition formed a felsic upper crust and a residual, mafic lower crust.

3.18.5.2.2

Arc thickening, intracrustal differentiation, and mixing

However, fractionation of primitive andesite is clearly just one of many possibilities for the genesis of continental crust with the composition of high Mg# andesite. An alternative view is that intracrustal differentiation, perhaps involving residual garnet and/or magnetite, has converted

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Geochemistry of Subduction-related Magmatic Arcs basaltic arc crust to andesitic continental crust. One mechanism for this is suggested by intermediate to felsic, garnet “cumulates” in the Talkeetna lower crust. These may have been derived from intermediate plutonic rocks and/or evolved volcanics that were gradually buried in the growing arc edifice (Figure 26, also see figure 9 in Kuno (1968)). Later, heating of the lower crust formed light REE enriched, heavy REE depleted melts that were extracted from these rocks. Mixing between primitive basalt and crustal melts with garnet-bearing residues could form high Mg# andesite (although we see little, if any, evidence for such mixing among Talkeetna arc volcanics). Then, delamination of the residues of lower crustal melting could yield an andesitic bulk composition for the entire crust.

3.18.6 CONCLUSIONS

Figure 26 Schematic illustration of the process of progressive burial of early formed plutonic and volcanic rocks within a growing arc edifice (also see Kuno, 1968, figure 9). Pyroxenites near the base of the crust are always denser than the underlying mantle, and temperatures are high near the Moho, so that these ultramafic cumulates may delaminate repeatedly whenever their thickness exceeds some critical value (e.g., Jull and Kelemen, 2001). Increasing pressure forms abundant garnet in Alrich, mafic, gabbroic rocks near the base of the section, and these too may delaminate. Intermediate to felsic plutonic rocks, and even volcanics, may be buried to lower crustal depths, where they undergo partial melting. Mixing of lower crustal melts with primitive basalt could produce high Mg# andesite magmas. This last process, together with delamination, provides a possible explanation for the formation of andesitic continental crust from initially basaltic arc crust.

Arc magmas are dramatically different, on average, from MORBs. Together with boninites, primitive and high Mg# andesites exemplify these differences. Unfortunately, primitive andesites are rare in intra-oceanic arcs, and so there is some possibility that they are produced by special intracrustal differentiation processes specific to continental arcs. However, there are primitive andesites in the oceanic Aleutian arc, and these lack any evidence for crustal assimilation, or even for recycling of components from subducted, continental sediments. Aleutian primitive andesites, and other primitive andesites with high Sr/Nd, cannot be produced by the mixing of primitive arc basalts and lower crustal melts. In most respects, the Aleutian primitive andesites are similar to all other primitive andesites, indicating that these magmas can be derived from primary, andesite melts. Primary andesites are probably produced by interaction between mantle peridotite and a partial melt of eclogite facies, subducting sediment and/or basalt. Thermal modeling cannot be used to rule out partial melting of subducting material beneath arcs, as has previously been supposed. Some characteristics of all primitive arc magmas, such as enrichment in thorium and lanthanum, and large depletions of niobium and tantalum relative to thorium and lanthanum, are best explained by a similar process, in which thorium and lanthanum are carried into the mantle wedge in a partial melt of subducting material in eclogite facies. Thus, the difference between primitive basalt and andesite may be largely determined by the relative proportions of mantle versus subduction zone melt components, together with a fluid component derived from relatively shallow dehydration of subducting materials.

References Although we favor the notion that fractionation of olivine and/or pyroxene from primitive andesites has played a key role in the formation of continental crust, which has a high Mg# andesite composition, this process alone is not sufficient to change the Mg# of the crust from a value of ,70, typical for primary melts in equilibrium with mantle olivine, to the value of ,50 estimated for bulk continental crust. Instead, removal of the crystalline products of fractionation from the crust into the mantle is also required. In fact, all “recipes” for continental crust probably require delamination. We show that the Talkeetna and Kohistan arc sections provide evidence for such a process, in the sense that they only contain a small proportion of the primitive, lower crustal cumulates that must have originally formed. The remaining primitive cumulates, which are pyroxenites and garnet granulites, were denser than the underlying mantle during arc crustal formation, while overlying gabbronorites were less dense. Thus, the Mg# gap in both sections is observed only where rocks become denser than residual peridotite. Removal of pyroxenite and garnet granulite from the base of arc crust leaves remaining crust with lower Sr/Nd and Pb/Ce than in primary arc magmas. Delaminated, recycled primitive arc cumulates would evolve to isotope signatures as depleted as, or more depleted than, the MORB source.

ACKNOWLEDGMENTS We thank Sue DeBari for advice and support for Andrew Greene, Ba¨rbel Sarbas for information from the GeoRoc database, Nate Hart for help with data entry, John Eiler for discussions, Tim Grove, Simon Turner, and Terry Plank for sharing preprints and sane advice, and Roberta Rudnick for forbearance during our long struggle to complete this chapter. Formal and informal reviews of a draft manuscript by Terry Plank, Steve Eggins, Simon Turner, Chris Hawkesworth, Bernard Bourdon, and Charlie Langmuir were both helpful and tolerant. We hope they find the finished product substantially improved. Research by Kelemen and Hanghøj for this chapter was supported in part by National Science Foundation Research Grants EAR-0125919, EAR-0087706, EAR-9910899, and OCE-9819666, plus the Charles Francis Adams Senior Scientist Chair at Woods Hole Oceanographic Institution.

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q 2003, Published by Elsevier Ltd. No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without prior written permission of the Publisher.

Treatise on Geochemistry ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp. 593–659

APPENDIX 1.

Periodic Table of the Elements.

197

198

Appendix APPENDIX 2.

a

APPENDIX 2. (Continued).

Table of Isotopes .

A

Element

40

Ar

99.600%

39 40

K K

41

K

93.2581% 1.277eþ9 yr 0.0117% 6.7302%

40 41 42 43 44 46 48

Ca Ca Ca Ca Ca Ca Ca

45

Sc

96.941% 1.03eþ5 yr 0.647% 0.135% 2.086% 0.004% 6eþ18 yr 0.187% 100%

46 47 48 49 50

Ti Ti Ti Ti Ti

8.0% 7.3% 73.8% 5.5% 5.4%

99.762% 0.038% 0.200%

50

V

51

V

1.4eþ17 yr 0.250% 99.750%

F

100%

50

Cr

20 21 22

Ne Ne Ne

90.48% 0.27% 9.25%

22 23

Na Na

2.6019 yr 100%

51 52 53 54

Cr Cr Cr Cr

1.8eþ17 yr 4.345% 27.7025 d 83.789% 9.501% 2.365%

24 25 26

Mg Mg Mg

78.99% 10.00% 11.01%

53 54 55

Mn Mn Mn

3.74eþ6 yr 312.3 d 100%

26 27

Al Al

7.17eþ5 yr 100%

28 29 30 32

Si Si Si Si

92.23% 4.67% 3.10% 150 yr

54 56 57 58 60

Fe Fe Fe Fe Fe

5.8% 91.72% 2.2% 0.28% 1.5eþ6 yr

E

C

59 60

Co Co

100 5.2714 yr

B

31 32 33

P P P

100% 14.262 d 25.34 d

C C

32 33 34 35 36

S S S S S

95.02% 0.75% 4.21% 87.32 d 0.02%

58 59 60 61 62

Ni Ni Ni Ni Ni

68.077% 7.6eþ4 yr 26.223% 1.140% 3.634%

35 36 37

Cl Cl Cl

75.77% 3.01eþ5 yr 24.23%

63 65 64 65 66 67 68

Cu Cu Zn Zn Zn Zn Zn

69.17% 30.83% 48.6% 244.26 d 27.9% 4.1% 18.8%

36 37 38

Ar Ar Ar

0.337% 35.04 d 0.063%

69 71

Ga Ga

60.108% 39.892%

39

Ar

269 yr

70

Ge

21.23%

A

Element

Abundance/half-life

1 2 3

H H H

99.985% 0.015% 12.33 yr

3 4

He He

0.000137% 99.999863%

6 7

Li Li

7.5% 92.5%

7 9 10

Be Be Be

53.12 d 100% 1.51eþ6 yr

10 11

B B

19.9% 80.1%

12 13 14

C C C

98.90% 1.10% 5,730 yr

14 15

N N

99.634% 0.366%

16 17 18

O O O

19

Source

C, B

C C

C, B

C

C

C

C, B

C (Continued )

Abundance/half-life

Source

C

B

E C

C

B

(Continued )

Appendix APPENDIX 2. (Continued).

APPENDIX 2. (Continued). A

Element

Abundance/half-life

72 73 74 76

Ge Ge Ge Ge

27.66% 7.73% 35.94% 7.44%

75

As

100%

74 76 77 78

Se Se Se Se

0.89% 9.36% 7.63% 23.78%

79 81

Br Br

50.69% 49.31%

78 80 81 82 83 84 85

Kr Kr Kr Kr Kr Kr Kr

0.35% 2.25% 2.29eþ5 yr 11.6% 11.5% 57.0% 10.756 yr

85 87

Rb Rb

72.165% 4.75eþ10 yr 27.835%

84 86 87 88 90

Sr Sr Sr Sr Sr

0.56% 9.86% 7.00% 82.58% 28.79 yr

89

Y

100%

90 91 92 93 94 96

Zr Zr Zr Zr Zr Zr

51.45% 11.22% 17.15% 1.53eþ6 yr 17.38% 3.8eþ19 yr 2.80 2%

93

Nb

100%

92 94 95 96 97 98 100

Mo Mo Mo Mo Mo Mo Mo

14.84% 9.25% 15.92% 16.68% 9.55% 24.13% 1.00eþ19 yr 9.63%

99

Tc

2.111eþ5 yr

98 99 100 101 102 104

Ru Ru Ru Ru Ru Ru

1.88% 12.7% 12.6% 17.0% 31.6% 18.7%

103 Rh

100%

102 Pd

1.02%

199

Source

C

B

B

A

Element

104 105 106 107 108 110

Pd Pd Pd Pd Pd Pd

51.839% 48.161%

106 108 110 111 112 113

1.25% 0.89% 12.49% 12.80% 24.13% 7.7eþ15 yr 12.22% 28.73%

Cd Cd Cd Cd Cd Cd

114 Cd 113 In 115 In

4.3% 4.41eþ14 yr 95.7%

112 114 115 116 117 118 119 120 122 124

Sn Sn Sn Sn Sn Sn Sn Sn Sn Sn

0.97% 0.65% 0.34% 14.53% 7.68% 24.23% 8.59% 32.59% 4.63% 5.79%

121 Sb 123 Sb

57.36% 42.64%

120 Te 122 Te 123 Te

0.096% 2.603% 1eþ13 yr 0.908% 4.816% 7.139% 18.95% 2.2eþ24 yr 31.69%

Te Te Te Te

130 Te

7.9eþ20 yr 33.80%

127 I 129 I

100% 1.57eþ7 yr

124 Xe

1.6eþ14 yr 0.10% 0.09% 1.91% 26.4% 4.1% 21.2% 26.9% 10.4%

126 128 129 130 131 132 134

Xe Xe Xe Xe Xe Xe Xe

136 Xe (Continued )

11.14% 22.33% 27.33% 6.5eþ6 yr 26.46% 11.72%

107 Ag 109 Ag

124 125 126 128

E

Abundance/half-life

Source

E

E, C, B

2.36e þ 21 yr 8.9% (Continued )

200

Appendix APPENDIX 2. (Continued).

APPENDIX 2. (Continued). A

Element

Abundance/half-life

133 Cs 134 Cs 137 Cs

100% 2.0648 yr 30.07 yr

130 132 134 135 136 137 138

0.106% 0.101% 2.417% 6.592% 7.854% 11.23% 71.70%

Ba Ba Ba Ba Ba Ba Ba

138 La 139 La

Source

B B

1.05eþ11 yr 0.0902% 99.9098%

A

Element

162 Dy 163 Dy 164 Dy

25.5% 24.9% 28.2%

165 Ho

100%

162 164 166 167 168 170

0.14% 1.61% 33.6% 22.95% 26.8% 14.9%

Er Er Er Er Er Er

169 Tm

100%

168 171 172 173 174 176

3.05% 14.3% 21.9% 16.12% 31.8% 12.7%

Yb Yb Yb Yb Yb Yb

138 Ce 140 Ce 142 Ce

0.25% 88.48% 5eþ16 yr 11.08%

141 Pr

100%

142 Nd 143 Nd 144 Nd

27.13% 12.18% 2.29eþ15 yr 23.80% 8.30% 17.19% 5.76% 1.1eþ19 yr 5.64%

176 177 178 179 180

no stable or long-lived isotope

180 Ta

145 146 148 150

Nd Nd Nd Nd

–

Pm

144 Sm 146 Sm 147 Sm

150 Sm 152 Sm 154 Sm

3.1% 1.03eþ8 yr 1.06eþ11 yr 15.0% 7eþ15 yr 11.3% 2eþ15 yr 13.8% 7.4% 26.7% 22.7%

151 Eu 153 Eu

47.8% 52.2%

152 Gd

1.08eþ14 yr 0.20% 2.18% 14.80% 20.47% 15.65% 24.84% 21.86%

148 Sm 149 Sm

154 155 156 157 158 160

Gd Gd Gd Gd Gd Gd

159 Tb

100%

156 158 160 161

0.06% 0.10% 2.34% 18.9%

Dy Dy Dy Dy

E

175 Lu 176 Lu

97.41% 3.78eþ10 yr 2.59%

174 Hf

2.0eþ15 yr 0.162% 5.206% 18.606% 27.297% 13.629% 35.100%

Hf Hf Hf Hf Hf

181 Ta 180 W 182 W 183 W 184 W 186 W

0.13% 26.3% 1.1eþ17 yr 14.3% 3eþ17 yr 30.67% 28.6% 37.40% 4.35eþ10 yr 62.60%

184 Os

5.6eþ13 yr 0.02% 2.0eþ15 yr 1.58% 1.6% 13.3% 16.1% 26.4% 41.0%

187 188 189 190 192

Os Os Os Os Os

Source

1.2eþ15 yr 0.012% 99.988%

185 Re 187 Re

186 Os

(Continued )

Abundance/half-life

191 Ir 193 Ir

37.3% 62.7%

190 Pt

6.5eþ11 yr 0.01% (Continued )

Appendix APPENDIX 2. (Continued).

APPENDIX 2. (Continued). A

Element

192 194 195 196 198

Pt Pt Pt Pt Pt

Abundance/half-life

Source

0.79% 32.9% 33.8% 25.3% 7.2%

A

Element

Abundance/half-life

Source Th232 U238

234 Th

1.9116 yr 7.538eþ4 yr 1.405eþ10 yr 100% 24.10 d

231 Pa 234 Pa

32760 yr 6.70 h

U235 U238

234 U

2.455eþ5 yr 0.0055% 7.038eþ8 yr 0.7200% 4.468eþ9 yr 99.2745%

228 Th 230 Th 232 Th

197 Au

100%

196 198 199 200 201 202 204

Hg Hg Hg Hg Hg Hg Hg

0.15% 9.97% 16.87% 23.10% 13.18% 29.86% 6.87%

203 205 206 207 208 210

Tl Tl Tl Tl Tl Tl

29.524% 70.476% 4.199 min 4.77 min 3.053 min 1.30 min

U238 U235 Th232 U238

U238 U235 Th232 U238

204 Pb

201

235 U 238 U

207 208 210 211 212 214

Pb Pb Pb Pb Pb Pb

1.4eþ17 yr 1.4% 22.1% 52.4% 22.3 yr 36.1 min 10.64 h 26.8 min

209 210 211 212 214 215

Bi Bi Bi Bi Bi Bi

100% 5.013 d 2.14 min 60.55 min 19.9 min 7.6 min

U238 U235 Th232 U238 U235

210 211 212 214 215 216 218

Po Po Po Po Po Po Po

138.376 d 0.516 s 0.299 ms 164.3 ms 1.781 ms 0.145 s 3.10 min

U238 U235 Th232 U238 U235 Th232 U238

215 At 218 At

0.10 ms 1.5 s

U235 U238

219 Rn 220 Rn 222 Rn

3.96 s 55.6 s 3.8235 d

U235 Th232 U238

223 Fr

21.8 min

U235

223 224 226 228

Ra Ra Ra Ra

11.435 d 3.66 d 1600 yr 5.75 yr

U235 Th232 U238 Th232

227 Ac 228 Ac

21.773 yr 6.15 h

U235 Th232 (Continued )

U238

Sources of short-lived radionuclides: B, bomb or reactor sources; C, cosmogenic; E, extinct radioactivities; U235, U238, Th232—nuclides in respective decay chains. Note: the symbol e indicates that the number following is that raised to the power of 10. a Modified from: Lawrence Berkeley Laboratory web site: http://ie.lbl.gov/education/isotopes.htm

202

Appendix APPENDIX 3. The Geologic Timescale.

Eon

Era

Period

Phanerozoic

Epoch

Millions of years ago

Holocene (Quaternary)

0.011 Pleistocene 1.82

Cenozoic

Pliocene 5.32 Miocene 23 (Tertiary)

Oligocene 33.7 Eocene 55 Paleocene 65

Cretaceous 144 Mesozoic

Jurassic 200 Triassic 250 Permian 295 Carboniferous Pennsylvanian 320 Mississippian 355

Paleozoic

Devonian 410 Silurian 440 Ordovician 500 Cambrian 543

Proterozoic Archean

Oldest rock Age of the solar system

2,500 4,400 4,550

Appendix APPENDIX 4.

Useful Values.

Molecular mass of dry air, ma ¼ 28:966 Molecular mass of water, mw ¼ 18:016 Universal gas constant, R ¼ 8:31436 J mol21 K21 Gas constant for dry air, Ra ¼ R=ma ¼ 287:04 J kg21 K21 Gas constant for water vapor, Rv ¼ R=mw ¼ 461:50 J kg21 K21 Molecular weight ratio 1 ; mw =ma ¼ Ra =Rv ¼ 0:62197 Stefan’s constant s ¼ 5:67 £ 1028 W m22 K24 Acceleration due to gravity, g (m s22) as a function of latitude w and height z (m) g ¼ ð9:78032 þ 0:005172 sin2 w 2 0:00006 sin2 2wÞð1 þ z=aÞ22 Ð Mean surface value, g ¼ 0p=2 g cosw dw ¼ 9:7976 Radius of sphere having the same volume as the Earth, a ¼ 6,371 km (equatorial radius ¼ 6,378 km, polar radius ¼ 6,357 km) Rotation rate of Earth, V ¼ 7:292 £ 1025 s21 Mass of Earth ¼ 5.977 £ 1024 kg Mass of atmosphere ¼ 5.3 £ 1018 kg Mass of ocean ¼ 1400 £ 1018 kg Mass of groundwater ¼ 15.3 £ 1018 kg Mass of ice caps and glaciers ¼ 43.4 £ 1018 kg Mass of water in lakes and rivers ¼ 0.1267 £ 1018 kg Mass of water vapor in atmosphere ¼ 0.0155 £ 1018 kg Area of Earth ¼ 5.10 £ 1014 m2 Area of ocean ¼ 3.61 £ 1014 m2 Area of land ¼ 1.49 £ 1014 m2 Area of ice sheets and glaciers ¼ 1.62 £ 1013 m2 Area of sea ice ¼ 1.9 £ 1013 m2 in March and 2.9 £ 1013 m2 in September (averaged between 1979 and 1987)

203