Journal of Petrology Advance Access originally published online on March 17, 2006
Journal of Petrology 2006 47(6):1051-1093; doi:10.1093/petrology/egl002
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A Detailed Geochemical Study of Island Arc Crust: the Talkeetna Arc Section, South–Central Alaska
1 DEPARTMENT OF EARTH AND OCEAN SCIENCES, UNIVERSITY OF BRITISH COLUMBIA, VANCOUVER, BC V6T 1Z4, CANADA
2 DEPARTMENT OF GEOLOGY, WESTERN WASHINGTON UNIVERSITY, BELLINGHAM, WA 98225, USA
3 LAMONT–DOHERTY EARTH OBSERVATORY, COLUMBIA UNIVERSITY, PALISADES, NY 10964, USA
4 DEPARTMENT OF GEOLOGY AND GEOPHYSICS, WOODS HOLE OCEANOGRAPHIC INSTITUTION, WOODS HOLE, MA 02543, USA
5 DEPARTMENT OF GEOLOGY AND PETROLEUM GEOLOGY, UNIVERSITY OF ABERDEEN, ABERDEEN AB24 3UE, UK
RECEIVED OCTOBER 10, 2003; ACCEPTED JANUARY 10, 2006
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ABSTRACT |
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The Early to Middle Jurassic Talkeetna Arc section exposed in the Chugach Mountains of south–central Alaska is 5–18 km wide and extends for over 150 km. This accreted island arc includes exposures of upper mantle to volcanic upper crust. The section comprises six lithological units, in order of decreasing depth: (1) residual upper mantle harzburgite (with lesser proportions of dunite); (2) pyroxenite; (3) basal gabbronorite; (4) lower crustal gabbronorite; (5) mid-crustal plutonic rocks; (6) volcanic rocks. The pyroxenites overlie residual mantle peridotite, with some interfingering of the two along the contact. The basal gabbronorite overlies pyroxenite, again with some interfingering of the two units along their contact. Lower crustal gabbronorite (
10 km thick) includes abundant rocks with well-developed modal layering. The mid-crustal plutonic rocks include a heterogeneous assemblage of gabbroic rocks, dioritic to tonalitic rocks (30–40% area), and concentrations of mafic dikes and chilled mafic inclusions. The volcanic rocks (
7 km thick) range from basalt to rhyolite. Many of the evolved volcanic compositions are a result of fractional crystallization processes whose cumulate products are directly observable in the lower crustal gabbronorites. For example, Ti and Eu enrichments in lower crustal gabbronorites are mirrored by Ti and Eu depletions in evolved volcanic rocks. In addition, calculated parental liquids from ion microprobe analyses of clinopyroxene in lower crustal gabbronorites indicate that the clinopyroxenes crystallized in equilibrium with liquids whose compositions were the same as those of the volcanic rocks. The compositional variation of the main series of volcanic and chilled mafic rocks can be modeled through fractionation of observed phase compositions and phase proportions in lower crustal gabbronorite (i.e. cumulates). Primary, mantle-derived melts in the Talkeetna Arc underwent fractionation of pyroxenite at the base of the crust. Our calculations suggest that more than 25 wt % of the primary melts crystallized as pyroxenites at the base of the crust. The discrepancy between the observed proportion of pyroxenites (less than 5% of the arc section) and the proportion required by crystal fractionation modeling (more than 25%) may be best understood as the result of gravitational instability, with dense ultramafic cumulates, probably together with dense garnet granulites, foundering into the underlying mantle during the time when the Talkeetna Arc was magmatically active, or in the initial phases of slow cooling (and sub-solidus garnet growth) immediately after the cessation of arc activity. KEY WORDS: island arc crust; layered gabbro; Alaska geology; island arc magmatism; lower crust
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INTRODUCTION |
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A major obstacle to our understanding the sources of island arc magmas is the effect of crustal evolution on the erupted volcanic products (Leeman, 1983
). Studies of volcanic rocks in active island arcs often suggest the existence of large sections of complementary cumulates in the lower crust (e.g. Kay & Kay, 1985). To improve our understanding of the lower crust in island arcs, studies often use plutonic xenoliths (e.g. Arculus & Wills, 1980; Conrad et al., 1983; DeBari et al., 1987), partial crustal exposures (e.g. Kawate & Arima, 1998; DeBari et al., 1999; Spandler et al., 2003), or seismic velocity measurements (e.g. Suyehiro et al., 1996; Fleidner & Klemperer, 1999; Holbrook et al., 1999). These provide indirect evidence of the magmatic processes in the deep crust of island arcs.
Exposures of island arc crust provide an opportunity to directly observe relationships not normally observable in active systems. Unfortunately, there are few examples of well-exposed island arc crust [e.g. Kohistan section, Pakistan (Tahirkheli, 1979); Darb Zubaydah, Saudi Arabia (Quick, 1990); Hokkaido section, Japan (Takashima et al., 2002)]. The Talkeetna Arc section in south–central Alaska has been recognized as the crust and upper mantle of an accreted, Early to Middle Jurassic island arc (Burns, 1985; DeBari & Coleman, 1989; Plafker et al., 1989) and hence offers a rare opportunity to directly compare volcanic rocks and their complementary deep and middle crustal plutonic equivalents.
The Talkeetna Arc has a fairly consistent north-dipping stratigraphy along strike with deepest rocks in the south and shallower rocks in the north (Fig. 1). The deepest level of the arc consists of residual mantle peridotite, overlain by pyroxenite, in turn overlain by gabbronorite. The lower crust is dominated by layered gabbronorite (10 km thick). Mid-crustal exposures are a heterogeneous assemblage of dioritic to tonalitic rocks mixed with gabbroic rocks and areas of abundant mafic dikes and chilled mafic inclusions. The upper crust of the arc is composed of thick sequences of lavas, tuffs and volcaniclastic debris-flow deposits of the Talkeetna Volcanic Formation (7 km thick) ranging from basalt to rhyolite.
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This study examines a diverse suite of Talkeetna Arc rocks from the Chugach and Talkeetna Mountains using mineral and whole-rock chemistry to assess the nature of the relationship between the cumulate gabbronorite and volcanic and upper-level plutonic rocks. These observations have been used to model the magmatic processes that link the volcanic and plutonic complements of this island arc and to place constraints on the nature of parental magmas.
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DESCRIPTION OF THE ARC |
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The Talkeetna Arc represents the northernmost exposure of an island arc system that may have extended continuously along the entire western margin of North America during the Early to Middle Jurassic (Plafker et al., 1989
). The arc is part of the Peninsular Terrane, which formed the overriding plate during subduction of the oceanic Farallon Plate (Plafker et al., 1989). The Peninsular Terrane has been thrust to the south over the accretionary rocks of the Chugach Terrane.
Early to Middle Jurassic plutonic and volcanic rocks of the Talkeetna Arc extend for more than 1000 km across southern Alaska. This study looks at a 150 km segment in the eastern portion of the arc in the Chugach Mountains (Fig. 1), but includes references to parts of the arc in the Talkeetna Mountains and on the Alaska Peninsula. Arc activity may have initiated between 210 and 200 Ma and continued in an oceanic setting until at least 180 Ma, whereafter magmatism may have shifted northward (Clift et al., 2005).
The earliest ages for the volcanic deposits from the Talkeetna Arc are 200 Ma (Pálfy et al., 1999) and 207 ± 3 Ma (J. Amato, personal communication, 2004) based on zircons. The Tuxedni Formation, unconformably overlying the Talkeetna Volcanics in the Talkeetna Mountains, contains Early Bajocian molluscs dated at 172 Ma (Imlay & Detterman, 1973; Imlay, 1984; Hillebrandt et al., 1992). Thus the volcanic rocks constrain the age of the arc to be between 207 and 172 Ma. Zircon ages from plutonic rocks in the Chugach Mountains generally agree with these constraints (201 to 181 Ma), but extend to younger ages for plutonic rocks further west on the Alaska Peninsula (183–164 Ma) and further north in the Talkeetna Mountains (177–156 Ma) (Rioux et al., 2001, 2002, 2004).
The buoyant arc crust was incorporated into the North American Plate by the Late Jurassic or Middle Cretaceous (Plafker et al., 1989). Up to 1000 km of Cenozoic strike-slip displacement may have transported the Talkeetna Arc to its present position (Plafker et al., 1989).
Previous studies and our continuing work in the eastern part of the arc have not identified an older, pre-arc, felsic crustal component within the section (Martin et al., 1915; Grantz et al., 1963; Detterman & Harstock, 1966; Newberry et al., 1986; Millholland et al., 1987; Plafker et al., 1989; Nokleberg et al., 1994; Rioux et al., 2001, 2002, 2004). For this reason, the Talkeetna section is interpreted as having formed in an island arc within oceanic crust, as distinct from an arc emplaced within pre-existing continental crust. In contrast, the western part of the arc (not part of this study) may have intruded into pre-existing Paleozoic basement of the Peninsular Terrane. In this western region there are some (detrital?) zircons in volcanic rocks on the Alaska Peninsula (Pálfy et al.,1999) and possible inheritance in zircons from Jurassic plutonic rocks on Kodiak Island (Roeske et al., 1989).
Most of the rocks examined for this study are exposed in the Chugach Mountains between Scarp Mountain and the Matanuska Glacier (Fig. 1), with additional volcanic rocks from the southern Talkeetna Mountains to the NW. The exposures extend over an area 5–18 km wide (perpendicular to the Border Ranges Fault) and nearly 150 km long (parallel to the Border Ranges Fault). The rocks analyzed in this study comprise four lithological units, in order of decreasing depth: (1) basal gabbronorite exposed on Scarp and Bernard Mountains that directly overlies mantle rock (Fig. 1); (2) lower crustal gabbronorites exposed between Tazlina Lake and Barnette Creek; (3) mid-crustal plutonic rocks primarily between the Nelchina and Matanuska Glaciers that consist of gabbronorite, gabbro, amphibole gabbronorite, diorite, quartz diorite, tonalite and very minor amounts of granodiorite, all with dikes and inclusions of chilled mafic rocks; (4) volcanic rocks from the length of the study area between Scarp Mountain and the Matanuska Glacier, plus a small group from the Talkeetna Mountains.
The boundary between lower crustal gabbronorite and mid-crustal plutonic rocks is inferred, from geochemistry and field relationships, to be between the extensive section of layered gabbronorites showing distict modal layering (lower crust) and the heterogeneous assemblages of mafic to felsic plutonic rocks (mid-crust). In addition, high-level amphibole gabbronorite forms a homogeneous pluton closely associated with volcanic rocks on Pippin Ridge, and an intrusion into volcanic rocks on Sheep Mountain.
Ultramafic rocks, not analyzed as part of this study, are exposed on four isolated hills (each 1–2 km2) at the eastern edge of the exposed arc section. In this locality, basal gabbronorite, garnet gabbro, pyroxenite and residual mantle peridotite form a layered crust–mantle boundary about 200 m thick separating outcrops with >90% gabbronorite from outcrops with >95% harzburgite and dunite (DeBari & Coleman, 1989).
South of these rocks, the base of the arc section is cut by the Border Ranges Fault that separates the arc rocks of the Peninsular Terrane from the accretionary sequences of the Chugach Terrane to the south (MacKevett & Plafker, 1974; Page et al., 1986). The Border Ranges Fault has a history as both a thrust and a right lateral strike-slip fault, which has been interpreted as the megathrust or backstop during accretion (Plafker et al., 1989).
Internally, the Talkeetna Arc section is disrupted by a network of arc-parallel faults and there is no continuous exposure from the base to upper crust. However, throughout the length of the arc section, the volcanics and mid-crustal plutonic rocks lie to the north of the lower crustal gabbronorites, with Moho and residual mantle sections even further to the south.
The arc crust also appears to have been tectonically thinned. Pressure estimates from garnet gabbros at the base of the crust are 850–1000 MPa, indicating a crustal thickness of 25–30 km (DeBari & Coleman, 1989; Mehl et al., 2001; Kelemen et al., 2003b). However, the exposed section has a maximum width of only 18 km, perpendicular to the strike of the Border Ranges Fault and to internal lithological contacts, which dip at steep to shallow angles to the north. Thus, the section has a present-day structural thickness of <18 km (Fig. 1).
The lithologies examined in this study and petrographic characteristics for 144 out of 159 plutonic rock samples collected are summarized in Table 1 and Fig. 2. Phase proportions in Table 1 are expressed as volume per cent, based on visual estimates. A representative suite of 83 volcanic rocks from the Talkeetna Arc was also analyzed for whole-rock chemistry (Clift et al., 2005).
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ANALYTICAL METHODS |
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Minerals in 21 samples (18 gabbronorites and three mafic dikes) were analyzed using a JEOL 733 electron microprobe at Massachusetts Institute of Technology and the University of Washington. Compositions of cores and rims in clinopyroxene, orthopyroxene, plagioclase, amphibole, spinel and olivine were measured with a 15 keV accelerating voltage and a 10 nA beam current. Element peaks/backgrounds were counted for 20–40 s and data were processed according to the Bence & Albee (1968)
matrix correction as modified by Albee & Ray (1970). Analytical error is <2% relative for major elements and <15% relative for trace elements. Averages for the analyses from each sample are listed in Table 2 (the full dataset is available for downloading from the Journal of Petrology website at http://www.petrology.oupjournals.org).
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Trace-element concentrations in clinopyroxene and plagioclase were measured using the Cameca IMS-3F ion microprobe at Woods Hole Oceanographic Institution (WHOI). Analytical techniques have been summarized by Shimizu & Hart (1982)
. A 20–30 nA primary beam of negative oxygen ions was focused to a 30 µm diameter. After preliminary sputtering, the emitted positive secondary ions were analyzed by a double focusing mass spectrometer using energy filtering with a secondary accelerating voltage of 4500 keV offset to –60 eV for the rare earth elements (REE) and –90 eV for other trace elements (Ti, V, Sr, Y, Zr). Analytical error is estimated as <10% relative to the concentrations. Averages for the analyses from each sample are shown in Table 3 (clinopyroxene REE concentrations), Table 4 (clinopyroxene trace-element concentrations), and Table 5 (plagioclase REE concentrations) (the full dataset is available at http://www.petrology.oupjournals.org).
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Whole-rock analyses have been acquired for 77 plutonic rocks from the Talkeetna Arc. Fifty-two analyses, with the major elements calculated on a volatile-free basis, are presented in Table 6. Samples were analyzed for 22 major and trace elements by X-ray fluorescence (XRF) and 27 trace elements by inductively-coupled plasma mass spectrometry (ICP-MS) at Washington State University (WSU) GeoAnalytical Laboratory. The preparation techniques and analytical methods for XRF (Johnson et al., 1999
) and ICP-MS (Knaack et al., 1994) are available from WSU GeoAnalytical Lab (http://www.wsu.edu/~geology/geolab/note/icpms.html).
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Nd isotopic ratios for 11 samples (six gabbronorites, three chilled mafic rocks, and two intermediate–felsic plutonic rocks) were measured at WHOI by MC-ICP-MS using a ThermoFinnigan Neptune system and are presented in Table 7. Between 50 and 300 mg of rock powders were spiked with Sm–Nd spike (enriched in 149Sm and 150Nd) and dissolved in a mixture 4 : 1 HF and HClO4. After drying, 6·2N HCl was added, samples were evaporated to dryness, and this was repeated. Then 0·5 ml of 2·5N HCl was added to the samples, they were left to stand for several hours, and were then transferred for column separation. The first columns separate light REE (LREE) [technique as described by Hart & Brooks (1977)
] and second columns separate Nd from Sm [technique as described by Richards et al. (1976)]. Samples were loaded, washed, and separated by time with 0·25N HCl for Nd and 0·6N HCl for Sm. Concentrations of Sm were determined by ICP-MS using a ThermoFinnigan Element system. Nd isotopic compositions were normalized to 146Nd/144Nd = 0·7219 and all results are corrected against the La Jolla Nd standard 143Nd/144Nd = 0·511847.
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MODAL MINERALOGY AND MINERAL CHEMISTRY OF THE GABBRONORITES |
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 TOPABSTRACTINTRODUCTIONDESCRIPTION OF THE ARCANALYTICAL METHODS MODAL MINERALOGY AND MINERAL... WHOLE-ROCK CHEMISTRYCALCULATED EQUILIBRIUM LIQUIDS...DISCUSSIONCONCLUSIONSUPPLEMENTARY DATAREFERENCES |
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Samples from lower crustal gabbronorites (between Tazlina Lake and Barnette Creek) were collected in north–south transects to characterize mineral compositions and evaluate vertical variations in mineral chemistry. Phase proportions (in weight per cent) for 16 of the gabbronorites were calculated using a least-squares method based on mass balance of whole-rock and mineral chemistry (Table 8).
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Clinopyroxene
Gabbronorites contain 1·80–22·0 wt % clinopyroxene, with an average of 13·4 wt % (Table 8). Clinopyroxene from lower crustal gabbronorite has a narrow compositional range. It is Mg-rich (13·1–15·7 wt % MgO) and shows minimal Fe enrichment with differentiation (Table 2). Magnesium number (Mg-number) [Mg-number = molar MgO/(MgO + FeOtotal) x 100, where FeOtotal refers to all Fe expressed as FeO] ranges from 73·6 to 81·3 (mean 77·3) and Al2O3 contents range from 1·32 to 3·24 wt % (Fig. 3). Clinopyroxene from basal gabbronorite (Scarp and Bernard Mountains) has high Al2O3 (4·69–7·31 wt %) compared with clinopyroxene in lower crustal gabbronorite with the same Mg-number. Clinopyroxene in amphibole gabbronorite, as part of the mid-crustal assemblage, has a range of Al2O3 contents similar to clinopyroxene in lower crustal gabbronorite, but with generally lower Mg-number (Pippin Ridge 69·2–70·7; Little Lake 71·5–75·4) except for gabbro intruding volcanics on Sheep Mountain (75·4–78·3) (Fig. 3a). TiO2 contents are <0·59 wt % and Na2O contents are <0·49 wt % in clinopyroxenes from all the gabbronorites.
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Clinopyroxene phenocrysts from two mafic dikes have Mg-numbers (72·1–80·5) similar to clinopyroxene in lower crustal gabbronorite (Fig. 3a). Alumina contents in the phenocrysts range from 2·11 to 6·08 wt %, and have greater core-to-rim variation than clinopyroxene in gabbronorite (Table 2). The phenocrysts have Cr2O3 contents as high as 0·67 wt %, whereas clinopyroxene in gabbronorite has <0·15 wt % Cr2O3 (except for two clinopyroxene analyses from basal gabbronorite sample 1710A4b that are higher).
Chondrite-normalized (N) REE patterns for clinopyroxene from lower crustal gabbronorite samples are parallel, with positive-sloping LREE segments [La/Sm(N) = 0·05–0·17; mean 0·11] and flat heavy REE (HREE) segments (5–25 x chondrite; mean 10 x chondrite; Fig. 4a). Ion microprobe analyses (3–8 per sample) showed minimal core-to-rim and grain-to-grain variation within samples, with a median standard deviation of less than 1 x chondrite. Rare earth element patterns for nine out of 13 lower crustal gabbronorite samples form a distinct band, with similar abundances. Clinopyroxene in four samples (mid-crustal amphibole gabbronorite 1709A2, 1723A5; lower crustal gabbronorite 1712A7, 1722A7) have noticeably higher REE abundances and negative Eu anomalies. Clinopyroxene in two basal gabbronorite samples (1710A4b, 1710A4d; dashed patterns in Fig. 4a) have LREE abundances that cross the band of patterns and have pronounced positive Eu anomalies. Clinopyroxene in one of these samples also has a positive Sr anomaly. Clinopyroxene phenocrysts in one mafic dike (1728A2) have similar major-element compositions and nearly identical REE abundances to clinopyroxene in lower crustal gabbronorites (Figs 3 and 4).
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In a chondrite-normalized trace-element diagram for clinopyroxene (Fig. 4b), Zr is depleted relative to LREE and Ti is depleted relative to the HREE. The clinopyroxene with higher REE abundances has negative Sr anomalies, as well as negative Eu anomalies. Four samples with the highest REE abundances in clinopyroxene are also enriched in Zr, Y, Ti, and V in clinopyroxene. Analyses of clinopyroxene from each of these samples form isolated trends in plots of Zr vs Y and Sr (Fig. 4c and 4d).
Orthopyroxene
Orthopyroxene compositions correlate with coexisting clinopyroxene compositions in the same samples, but have slightly lower Al2O3 (0·82–2·09 wt %) and Mg-number (Fig. 3b). The Mg-numbers for orthopyroxene are more variable than clinopyroxene Mg-numbers in lower crustal gabbronorite (65·0–74·9; mean 69·9), and are also lower for orthopyroxene in mid-crustal amphibole gabbronorite (Pippin Ridge 57·4–59·6; Little Lake 57·9–62·7; Sheep Mountain 65·6–65·8) (Table 2). TiO2 contents are <0·49 wt % and Na2O contents are <0·44 wt %. From mass-balance calculations, gabbronorites contain 1·71–45·7 wt % orthopyroxene (Table 8), averaging 13·4 wt %, the same as the average clinopyroxene proportion.
DeBari & Coleman (1989) estimated equilibrium conditions for basal gabbronorite of 800–925 °C at 0·95–1·05 GPa. Temperature estimates for coexisting pyroxenes for the suite of gabbronorites in this study, using the QUILF program (Andersen et al., 1993) with pressure set at 0·7 GPa, range from 740 to 930 °C, and at 0·3 GPa range from 700 to 920 °C. The calculated temperatures for pyroxenes from Barnette Creek were slightly lower than those from Tazlina Lake. Amphibole gabbronorite from Pippin Ridge, Little Lake, and Sheep Mountain yielded similar temperatures to those from Barnette Creek (800–820 °C).
Plagioclase
Plagioclase is the most abundant phase in the gabbronorites, ranging from 29 to nearly 80 wt % (Table 8). There is minimal intragrain zoning or variation in anorthite content. The anorthite content ranges from An91·8 to An77·7 in basal gabbronorite and An94·7 to An75·0 in lower crustal gabbronorite from Tazlina Lake and Barnette Creek (Table 2). There is a considerable gap between the plagioclase compositions in mid-crustal amphibole gabbronorite from Pippin Ridge and Little Lake (An65·5–An59·8) and the deeper gabbronorite. Most of the analyzed gabbronorites show reverse zoning in plagioclase, with rims of most grains slightly more calcic than the cores. Plagioclase REE patterns have negative-sloping LREE segments, with relatively low abundances, and pronounced positive Eu anomalies (Fig. 4e). Samples with the highest REE abundances (e.g. mid-crustal amphibole gabbronorite samples 1709A2, 1723A5) have the lowest An contents.
Amphibole
Amphibole in the gabbronorites constitutes between 0·97 and 50·1 wt % (mean 19·1 wt %) (Table 8). The amphibole is pargasitic to actinolitic hornblende, with some edenite component [nomenclature of Leake (1978)]. Mg-numbers for the amphibole range from 53·2 to 78·2, with slightly lower Mg-numbers in amphiboles from mid-crustal amphibole gabbronorites (Table 2). Mg-numbers in amphibole correlate with coexisting clinopyroxene Mg-numbers; Mg-numbers for clinopyroxene are mostly higher than those of amphibole, with the exception of one sample from Scarp Mountain (1710A4e). The amphibole in basal gabbronorites from Scarp Mountain also has higher Al2O3 than other amphibole, analogous to the pyroxene. As noted by Burns et al. (1991), amphibole in gabbronorite samples is interstitial, and so is interpreted as a late magmatic or high-temperature hydrothermal mineral.
Spinel
The proportion of Cr–Al–Mg–Fe–Ti oxide in gabbronorite ranges from 0·15 to 12·3 wt % (Table 8). With the exception of one sample from Scarp Mountain (1710A4d) that contains Fe–Ti oxide, spinels in basal gabbronorite are all Mg–Al pleonastes. All the other gabbronorites contain exsolved magnetite–ilmenite pairs. Lower crustal gabbronorites from Tazlina Lake average 10·2 wt % Fe–Ti oxide. TiO2 contents in magnetite from lower crustal gabbronorites average 3·30 wt % (Table 2).
Olivine
Olivine is present in only one gabbroic sample (Tazlina 1712A4). This sample has whole-rock chemistry similar to that of the other gabbronorites. The olivine grains are not zoned and are mostly resorbed. Thin sections reveal symplectite-like intergrowths of amphibole and magnetite between olivine and orthopyroxene. The olivine Mg-numbers in this sample range from 69·9 to 72·7 (Table 2).
Summary of mineral chemistry
The compositions of minerals from lower crustal gabbronorite between Tazlina Lake and Barnette Creek are consistent with a common igneous parentage, limited interaction with percolating evolved interstitial liquids, and/or re-equilibration of cores and rims through intracrystalline diffusion (Tribuzio et al., 1999). There is no discernible vertical variation (from north–south transects), nor is there any consistent core-to-rim variation in clinopyroxene, orthopyroxene or amphibole from any of the gabbronorites or chilled mafic rocks.
Plagioclase in the layered gabbronorite has distinctly high An contents (mean 86·1), characteristic of plagioclase crystallized from hydrous arc magmas (e.g. Sisson & Grove, 1993). In hydrous arc magmas, clinopyroxene crystallizes before plagioclase as a result of the suppression of plagioclase crystallization because of water dissolved in the melt. Thus, the most primitive clinopyroxene (high Mg-number, low Yb) will have negligible Eu anomalies, and more evolved clinopyroxene will have more pronounced negative Eu anomalies (Plank et al., 2004). This is reflected by the presence of pronounced positive Eu anomalies in plagioclase and the absence of negative Eu anomalies in clinopyroxene (in all but two lower crustal gabbronorites; see below). The rims of plagioclase in samples from Tazlina and Barnette are slightly more calcic than the cores, by an average of 1·5 mol % An. This reverse zoning may be the result of several factors, such as an increase of pH2O in interstitial liquid (Arculus & Wills, 1980) or diffusion of Na from plagioclase into late crystallizing amphibole (Khan et al., 1989).
Clinopyroxene may have acquired REE characteristics from subsolidus breakdown of plagioclase + olivine to form pyroxenes, spinel, and garnet at the base of the crust. As previously described, clinopyroxene in basal gabbronorite samples has positive Eu anomalies and flatter LREE segments than clinopyroxene in lower crustal gabbronorite, and one of these samples has a positive Sr anomaly (Fig. 4b). Positive Eu anomalies and elevated LREE in clinopyroxene from gabbro-derived granulites from the Northern Apennines have also been attributed to redistribution of REE during granulite-facies recrystallization (Montanini & Tribuzio, 2001).
The higher Al2O3 contents of pyroxene in the basal gabbronorite, compared with pyroxene in underlying pyroxenite, were attributed by DeBari & Coleman (1989) to Al increase in the liquid as a result of pyroxenite crystallization. Although we believe pyroxenite crystallization may have been extensive (see the pyroxenite fractionation section, below), Al contents in pyroxenes in basal gabbronorites are six to eight times higher than in underlying pyroxenite. Pyroxenite crystallization was probably not sufficient to cause a six- to eight-fold increase in incompatible element abundances. Thus, if the Al contents of pyroxene in basal gabbronorites and pyroxenites were both formed in equilibrium with liquid, this suggests the presence of both high- and low-Al melts entering the Talkeetna Arc crust through the Moho. Alternatively, the high Al content in pyroxene in basal gabbronorites may result from metamorphic reactions involving breakdown of plagioclase ± olivine. In addition to high-Al pyroxene in basal gabbronorite, high-Al clinopyroxenes are found in several mafic dikes (Fig. 3a). This suggests that either some high-Al pyroxene formed as igneous phases, or the high-Al pyroxene grains in the dikes are xenocrysts derived from granulite-facies rocks near the base of the arc crust.
Several lower crustal gabbronorite samples with anomalous clinopyroxene trace-element chemistry (Fig. 4) have a distinct mineralogy. Tazlina sample (1712A7) has <2 wt % clinopyroxene, 12·3 wt % magnetite and 40·9 wt % amphibole, and a sample from Barnette Creek (1722A7) has >45 wt % orthopyroxene (Table 8). The high proportion of mafic minerals may be responsible for the enrichment of Zr in clinopyroxene rims from these samples, as a result of the incompatibility of Zr in Fe–Ti oxides and amphibole (e.g. Tribuzio et al., 1999). Clinopyroxene rim compositions are enriched in Zr in mid-crustal amphibole gabbronorite. In this case, the rims probably crystallized from a more evolved magma and/or retained a greater proportion of trapped melt, enriched in incompatible elements. Low Sr contents and negative Eu anomalies in clinopyroxene from mid-crustal amphibole gabbronorite clearly reflect prior removal of plagioclase from the melt that crystallized the clinopyroxene (Fig. 4b).
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WHOLE-ROCK CHEMISTRY |
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Major- and trace-element compositions
Mid-crustal intermediate to felsic plutonic rocks from the Talkeetna Arc have remarkably similar major- and trace-element compositions to the volcanic rocks, whereas lower crustal ‘cumulate’ gabbronorites have systematically different major-element compositions, but many comparable trace-element characteristics. Major-element variations in whole-rocks are plotted in Fig. 5, along with previously published results for the Talkeetna Arc and published analyses from the Mesozoic Kohistan Island Arc in Pakistan. Talkeetna Arc samples have 38·0–77·9 wt % SiO2, with considerable overlap between the volcanic and intermediate to felsic plutonic rocks (Fig. 5). Al2O3, CaO, and MgO decrease with increasing SiO2. There is very little overlap in SiO2 between the gabbroic rocks (dominantly cumulates) and the volcanic and intermediate to felsic plutonic rocks. Al2O3 contents are high (mostly >15 wt % for samples with <65 wt % SiO2) for all lithologies; in particular, gabbroic rocks contain 12·3–28·1 wt % Al2O3 with a mean of 18·9 wt % (Fig. 5). TiO2 contents are low (primarily <1·5 wt %) for Talkeetna volcanic and intermediate to felsic plutonic rocks, with concentrations decreasing with increasing SiO2. Gabbroic rocks also define a broad, steeply decreasing trend of TiO2 with decreasing SiO2, at lower SiO2 contents. FeOtotal (not shown) behaves in the same way as TiO2 and defines a similar trend vs SiO2. There is minimal overlap in CaO and MgO between the gabbroic rocks and other lithologies, with most of the gabbroic rocks having higher abundances. Gabbroic rocks generally have lower K2O and Na2O (not shown) than the volcanic and intermediate to felsic plutonic rocks.
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Overall, compositions from the Talkeetna Arc for most of the major elements are similar to those for Kohistan Island Arc rocks, though many of the Kohistan samples have higher SiO2 at about the same Mg-number (Fig. 5e). A broad band of data points overlaps the tholeiitic and calc-alkaline fields in Mg-number vs SiO2 [using the classification scheme of Miyashiro (1974)
, expressed using Mg-number], with the majority of the basaltic volcanics lying within the tholeiitic field. Mg-numbers range from 38·1 to 77·4 (mean 57·2) for the gabbroic rocks and are mostly lower for the volcanic (14·5–62·6; mean 42·8) and intermediate to felsic plutonic rocks (16·4–67·6; mean 41·8). The chilled mafic rocks define steeply decreasing trends for both MgO and Mg-number vs SiO2, with MgO varying from 3·5 to 12·4 wt % (mean 7·2 wt %) and Mg-number from 39·7 to 69·8 (mean 56·3). Concentrations of Ni, V, and Sr are, on average, higher in cumulate gabbronorites than in volcanic and intermediate to felsic plutonic rocks (Fig. 5g–i), but similar to those in the chilled mafic rocks. Zr concentration increases from very low levels in cumulate gabbronorite to higher concentrations in volcanic and intermediate to felsic plutonic rocks, with a distinct group of volcanic rocks diverging from the main trend to even higher concentrations.
Chondrite-normalized REE patterns for the range of lithologies from the Talkeetna Arc (volcanic rocks, intermediate to felsic plutonic rocks, chilled mafic rocks, and gabbroic rocks) are remarkably parallel and increase in abundance systematically from the basal gabbronorite and lower crustal gabbronorite (1–10 x chondrite) through the volcanic upper crust (8–38 x chondrite) (Fig. 6a and c). The patterns are flat through the middle REE (MREE) and HREE, with very few crossing patterns. The gabbronorites with lower abundances are noticeably LREE depleted and have distinct positive Eu anomalies. The LREE segments progressively flatten with increasing REE abundance in gabbroic rocks. Patterns for the mid-crustal gabbros, and several chilled mafic rocks and intermediate to felsic plutonic rocks, are nearly flat with small Eu anomalies.
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The REE patterns for the volcanic samples throughout the arc section form a distinct band with parallel MREE to HREE segments. However, three lavas, out of the 84 analyzed, are strongly depleted in HREE, similar to a single HREE-depleted volcaniclastic sample reported by Plafker et al. (1989)
(Fig. 6a). The patterns in the main band are progressively LREE enriched with increasing abundance and samples with higher abundances generally have more pronounced negative Eu anomalies. REE patterns for volcanic and chilled mafic rocks found in close proximity (<1 km) are nearly identical. Several of the intermediate to felsic plutonic rocks are slightly LREE enriched and most of these patterns overlap volcanic rock REE patterns. Normal mid-ocean ridge basalt (N-MORB)-normalized trace-element patterns are characterized by high concentrations of large ion lithophile elements (LILE) Ba, K, and Pb and lower abundances of high field strength elements (HFSE) Nb, Ta, Zr, Hf, Ti, Y (Fig. 6b and d). The concentrations of virtually all incompatible elements rise incrementally, from basal gabbronorites through the upper volcanic rocks, with noticeably parallel patterns. Positive Pb and Sr spikes, relative to adjacent elements, are present for all lithologies, with the exception of eight volcanic rock samples with high trace-element concentrations, which have negative Sr and Eu anomalies. Ti concentrations are higher relative to HREE in cumulate gabbronorites from the Tazlina and Barnette areas, whereas Ti is depleted relative to HREE in most of the volcanic and intermediate to felsic plutonic rocks.
Summary of whole-rock chemistry
The major- and trace-element whole-rock chemistry for Talkeetna Arc rocks is consistent with evolution as part of a co-magmatic differentiation sequence related to similar parental magmas. The whole-rock patterns reflect variability that may be related primarily to fractional crystallization, where REE and HFSE increase and Mg-number decreases with differentiation. The representative suite of volcanic rocks spans the range of SiO2 contents, from 48–80 wt % SiO2. Sixty-three of the 83 volcanic rocks from throughout the arc have evolved compositions, with Mg-number <50.
The importance of plagioclase fractionation is reflected by variation in whole-rock major- and trace-element compositions. Plagioclase represents the dominant phase in the cumulate gabbronorites (Table 8) and trends of Al2O3 and CaO in the volcanic rocks are primarily the result of plagioclase fractionation. Positive Eu anomalies for layered gabbronorite clearly reflect their cumulate nature, assuming mantle-derived, parental arc magmas have no Eu anomaly prior to crystallization of plagioclase. Kelemen et al. (2003a) postulated that the high Pb and Sr in cumulate gabbroic rocks from the Talkeetna Arc result from a combination of (1) high plagioclase/liquid distribution coefficients for Pb and Sr, and (2) the presence of Pb and Sr spikes in parental Talkeetna Arc liquids.
Crystallization of Fe–Ti oxides in lower crustal gabbroic cumulates was an important control on the composition of the erupted arc lavas (see also, e.g. Osborn, 1959; Gill, 1981; Woodhead, 1988). Fe–Ti oxides represent a major phase within the lower and mid-crustal gabbronorites from the Talkeetna Arc, and clearly affected the whole-rock TiO2 budget for other lithologies in the arc crust. This is particularly evident in the trace-element patterns, where all of the Fe–Ti oxide-bearing lower crustal gabbronorite samples from the Tazlina–Barnette area have pronounced positive Ti anomalies, and nearly all of the volcanic and intermediate to felsic plutonic rocks have distinct negative Ti anomalies (Fig. 6). These complementary Ti anomalies imply that crystallization of Fe–Ti oxides within the layered gabbronorite caused liquids forming the volcanic and intermediate to felsic plutonic rocks to be depleted in TiO2. There probably was no Ti anomaly relative to HREE in primitive Talkeetna Arc magmas.
Yb can be used as a proxy for differentiation for the various arc lithologies given the systematic increase in concentrations from basal gabbronorite through the most evolved volcanics, and the flat HREE patterns in almost all samples. When plotted vs east–west distance along the Talkeetna Arc section [in meters, from Universal Transverse Mercator (UTM) coordinates], Yb(N) provides a means for visualizing the ‘chemical stratigraphy’ of the Talkeetna Arc crust (Fig. 7). The easternmost exposures represent the deepest portion of the arc crust, where basal gabbronorites overlying and interlayered with pyroxenite have the lowest whole-rock Yb(N), highest Mg-number, and high An contents. Most of the lower crustal gabbronorites from the Tazlina–Barnette area have similar chemical characteristics and show a slight sympathetic variation between Yb(N) and An, as expected for cumulates created during fractional crystallization. Mid-crustal amphibole gabbronorites from Pippin Ridge have higher Yb(N), lower An and Mg-number, and small Eu anomalies. The chilled mafic rocks and intermediate to felsic plutonic rocks clearly overlap the compositions of the nearby volcanic rocks for all elements, and probably represent liquid compositions (Kelemen et al., 2003a). The volcanic rocks extend to more evolved compositions in the eastern part of the arc section. However, intermediate to felsic plutonic rocks are either unexposed or missing in this area.
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Neodymium isotopic compositions
Nd isotopic compositions (143Nd/144Nd) for 11 plutonic samples of different lithologies (basal gabbronorite, lower crustal gabbronorite, chilled mafic rocks and intermediate to felsic plutonic rocks) range from 0·512960 to 0·513047 (Table 7). These samples have a narrow isotopic range and most of the 11 samples have the same ratio within analytical error. 143Nd/144Nd ratios for six volcanic samples (listed for comparison in Table 7) range from 0·512919 to 0·512998 (Clift et al., 2005
). 143Nd/144Nd ratios for additional plutonic samples in the Chugach Mountains (mid-crustal gabbronorite, and intermediate to felsic plutonic rocks) range from 0·512857 to 0·512988 (Rioux et al., 2004). Nd isotopic ratios from the Talkeetna Mountains have a slightly greater range (0·512815–0·513304) (Rioux et al., 2004).
Initial ratios were calculated at an age of 182 Ma. This age is based on the average of a range of U–Pb zircon ages from similar samples in close proximity in the Chugach Mountains (Rioux et al., 2001, 2002, 2004), and a Sm–Nd isochron age (the best-fitting isochron fit for eight of 11 samples) (Table 7).
As shown in Fig. 8, the initial Nd isotopic compositions are close to those of Jurassic Atlantic MORB (0·51276–0·51278) (Hoernle, 1998) and Jurassic Pacific MORB (0·51290–0·51301) (Hauff et al., 2003) and thus were derived from a mantle source very similar to the MORB source. The small difference between the initial Nd isotopic compositions in Talkeetna rocks and Jurassic MORB is similar to the small difference in Nd isotopes between Marianas arc lavas and present-day MORB.
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Two chilled mafic rocks (1728A3, 1719A11) have the highest Mg-numbers for noncumulate samples and represent the most primitive samples analyzed in the Talkeetna Arc. The fact that these samples have Nd isotopic compositions similar to the other lithologies indicates that assimilation of older crustal material with different Nd isotope ratios was minimal to nonexistent during the chemical differentiation of the Talkeetna Arc.
The Nd isotopic characteristics of the Talkeetna Volcanic Formation are consistent with an oceanic subduction-related origin, with no evidence of collision with North America during arc activity (Clift et al., 2005). Also, to a first-order approximation, the Nd isotopic data are consistent with derivation of the entire arc section from one type of primary magma derived from a mantle source with a fixed proportion of ‘subduction components’. However, continuing research on the Talkeetna Arc recently revealed slightly enriched Sr isotopic ratios in the younger plutonic rocks of the Talkeetna Mountains (north of the study area), which may reflect intrusion of rocks younger than 177 Ma into recently accreted continental material or into a pre-existing crustal boundary (Rioux et al., 2004).
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CALCULATED EQUILIBRIUM LIQUIDS FOR GABBRONORITES |
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TOPABSTRACTINTRODUCTIONDESCRIPTION OF THE ARCANALYTICAL METHODSMODAL MINERALOGY AND MINERAL...WHOLE-ROCK CHEMISTRYCALCULATED EQUILIBRIUM LIQUIDS...DISCUSSIONCONCLUSIONSUPPLEMENTARY DATAREFERENCES |
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Field and petrographic observations, and mineral and whole-rock chemistry described above, are strong evidence that the layered gabbronorites are cumulates that formed from partial crystallization of a magma from which the remaining liquid was subsequently removed. This remaining liquid was probably erupted as the basalts and basaltic andesites of the volcanic section. This assumption can be tested by determining whether the cumulate, layered gabbronorites crystallized in equilibrium with liquids that formed the volcanic rocks.
The trace-element composition of ‘equilibrium liquids’ can be determined by utilizing trace-element compositions of clinopyroxene from the gabbronorite cumulate rocks and clinopyroxene/liquid partition coefficients. However, this assumes that the clinopyroxenes from the gabbronorite cumulates have retained their igneous trace-element characteristics. Sixteen of the 18 gabbronorite samples have parallel, LREE-depleted trace-element patterns, and hence have indeed retained their igneous character. In contrast, clinopyroxenes from the basal gabbronorites have been re-equilibrated under granulite-facies conditions.
Liquid REE concentrations calculated to be in equilibrium with clinopyroxene in 18 gabbronorites and two mafic dikes are shown in Fig. 9a, along with a shaded area representing 80 of the 83 REE patterns for volcanic rocks throughout the arc (from Fig. 6a, omitting three anomalous LREE-enriched, HREE-depleted lavas). Calculated equilibrium liquids for all but two of the lower crustal gabbronorite samples between Tazlina Lake and Barnette Creek lie within the range of volcanic rock compositions [8 to 38 x chondrite; mean Ce/Yb(N) = 1·45; Fig. 9b]. Ten out of 14 calculated liquids from lower crustal gabbronorite samples form a band of REE patterns between 11 and 18 x chondrite [mean Ce/Yb(N) = 1·10]. A suite of eight chilled mafic rocks also ranges from 8 to 18 x chondrite with an average Ce/Yb(N) of 1·15.
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The above-described evidence from calculated liquids links many of the layered gabbronorites and volcanic rocks through crystal fractionation, such that the cumulate gabbronorites crystallized in equilibrium with liquids that were extracted and erupted to produce the volcanic rocks.
Some, but not all, of the lower crustal gabbronorite samples in apparent cpx/liquid REE exchange equilibrium with a given volcanic rock composition are also in cpx/liquid Fe/Mg exchange equilibrium with that same volcanic rock composition. However, 10 volcanic samples (Mg-number >54) have Mg-numbers too high for equilibrium with any of the lower crustal gabbronorites. This suggests that they formed in equilibrium with more primitive cumulates, either pyroxenites or gabbronorites that were not sampled during our study. It should be noted that Fe/Mg partitioning between clinoproxene and melt is linked to Fe3+/Fe2+ through fO2 and for calculations involving the layered gabbronorites, the values of Sisson & Grove (1993) were used: Fe/Mg Kdcpx/liquid of 0·23 calculated with total Feliquid. This assumption is a source of uncertainty, but it remains the case that for any reasonable combination of Fe3+/Fe2+ and Fe/Mg Kdcpx/liquid, about 10 of the volcanic samples have Mg-numbers too high for Fe/Mg equilibrium with lower crustal gabbronorite.
A plot of Yb(N) vs Mg-number shows that most of the volcanic rocks form a trend of increasing Yb(N) with decreasing Mg-number (Fig. 10), as expected for crystal fractionation. In contrast, clinopyroxenes in the lower crustal gabbronorites have nearly constant Mg-number and a wide range of Yb. Therefore, the calculated liquids in equilibrium with these clinopyroxenes do not show the same trend of Yb(N) vs Mg-number as the volcanic rocks.
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Interestingly, several of the volcanic rocks have a high Mg-number, high Yb signature, similar to calculated liquids for several gabbronorites (Fig. 10). The clinopyroxenes with this signature are from the mid-crustal amphibole gabbronorite and from orthopyroxene-rich (>40 wt %) sections of the lower crustal gabbronorite. It is possible that the high Mg-number, high Yb clinopyroxenes and high Mg-number, high Yb lavas represent complementary products of the same process. This signature could be the result of several processes: reaction of migrating, evolved melt ± aqueous fluids and mafic cumulates (e.g. Costa et al., 2002
), magma mixing, and/orfractionation of very oxide-rich cumulates. Mid-crustal amphibole gabbronorite samples have high REE abundances (patterns that overlap volcanic rock compositions), negative Sr and Eu anomalies in clinopyroxenes, and low An content in plagioclase. Clinopyroxene from mafic layers in the layered gabbronorite has high REE abundances, negative Sr and Eu anomalies, and higher concentrations of trace elements (Zr, Y, Ti, V). High Mg-number and high Yb values for these calculated liquids may be linked to the formation of the mafic layers. In the case of the lavas, the high Mg-number, high Yb samples could also be produced by mixing of evolved melt (low Mg-number, high Yb) and primitive melt (high Mg-number, low Yb). Mixed melts could then precipitate high Mg-number, high Yb cumulates. |
DISCUSSION |
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Magma compositions for the Talkeetna Arc
The geological evidence and chemical characteristics of the Talkeetna Arc are consistent with the hypothesis that the bulk of the crust formed during arc magmatism in an intra-oceanic setting. Almost all samples appear to be related to a single type of parental magma through processes of intracrustal differentiation. Most of the volcanic rocks have evolved compositions, with nearly 75% of the representative suite having >53 wt % SiO2 and Mg-number <50. Many of the evolved volcanic compositions are a result of cumulate processes recorded by middle to lower crustal gabbronorites.
One of our goals in this study is to determine the parental melt composition(s) for the cumulate gabbroic rocks and evolved lavas in the arc. This provides us with crucial information about magmatic fluxes into arc systems. However, this parental magma composition also provides us with starting points for two types of modeling that are described in subsequent sections of the paper: (1) crystal fractionation modeling, to quantify the processes and proportions of rock types involved in crustal differentiation; (2) pyroxenite addition models, to determine the composition of hypothetical primary arc melts in equilibrium with residual peridotite. To do this, we have assessed the most primitive magma compositions preserved in the Talkeetna Arc by selecting sampled primitive basaltic compositions (high Mg-number, Ni, Cr) (Table 6). In Table 6 we have only tabulated the volcanic samples with Mg-number >60, along with the mafic dikes and inclusions from this study. The complete set of volcanic rock analyses has been given by Clift et al. (2005).
As an overview, of the 83 volcanic rocks and eight chilled mafic rocks analyzed as a representative suite within the arc, six volcanic rocks and two chilled mafic rocks have Mg-number >60. Fourteen volcanic rocks have Mg-number 50–60. Of the 20 volcanic rocks with Mg-numbers between 50 and 70, SiO2 contents range from 47 to 76 wt % (with 11 having <53 wt % SiO2, two having 53–57 wt % SiO2, two having 57–63 wt % SiO2, and five having >63 wt % SiO2). All five of the eight chilled mafic rocks with Mg-number >50 are basaltic in composition.
In the process of surveying the suite of volcanic samples for potential parental melt compositions, several interesting observations arise. The volcanic samples with Mg-numbers >50 are not uniformly distributed throughout the study area. They are primarily concentrated in two areas, one in the easternmost exposures near Stuck Mountain and the other in the Little Oshetna River area, the most northerly of the sampled volcanic rocks, in the Talkeetna Mountains [see Clift et al. (2005) for descriptions]. Fifteen of the 20 volcanic samples with Mg-number >50 are from these two areas.
Five of the six lavas with Mg-number >60 are from these two areas and were not selected as potential parental melt compositions. The Little Oshetna River area is dominated by basaltic lavas, unconformably overlain by Middle Jurassic sandstone in close proximity (Clift et al., 2005). These lavas have trace-element compositions that are characteristic of eruption in an off-axis setting (Clift et al., 2005). Several lavas from Stuck Mountain have distinct trace-element chemistry (with high La/Yb; see discussion below) and were not used as parental compositions. The remaining lava, which has the highest Mg-number (62·6), has 57 wt % SiO2 and therefore it probably cannot be parental to evolved basalts and basaltic andesites. This eliminated all six lavas with Mg-number >60 as initial liquids for the crystal fractionation modeling.
The two primitive chilled mafic rocks with Mg-number >60 (1719A11, mafic pillow; 1728A3, mafic dike) are basaltic, with MgO contents of 8·35 and 8·09 wt %, Mg-numbers of 64·0 and 61·9, and Al2O3 contents of 18·5 and 19·3 wt %, respectively (Table 6). They also have high Cr concentrations (296 and 188 ppm) and variable Ni (205 and 73 ppm). The REE patterns for these primitive samples lie just below a field of calculated liquid REE for lower crustal gabbronorite, using the cpx/liquid coefficients of Hart & Dunn (1993). However, in detail the relationship of calculated liquids to actual liquids depends on the selection of partition coefficients, which is uncertain. The REE patterns for the chilled mafic pillow and a second, basaltic mafic pillow, with Mg-number 57·0, lie close to the calculated primary magma composition of DeBari & Sleep (1991). However, the REE patterns for these chilled mafic rocks have small positive Eu anomalies. Calculations from whole-rock and plagioclase REE concentrations measured on Talkeetna samples during this study indicate that the Eu anomalies can be accounted for by as little as 3 wt % accumulated plagioclase, with an insignificant effect on the major-element composition. In summary, we chose the two primitive chilled mafic rocks as potential compositions of primitive melts, parental to more evolved basalts and basaltic andesites, and perhaps to most andesites, dacites and rhyolites as well. Burns et al. (1991) also reported three mafic dikes from within the study area with 9·05–12·1 wt % MgO and Mg-numbers of 64·3–69·8. Two of these samples contain <14 wt % Al2O3, which may further indicate that they represent relatively unfractionated melts that quenched within the upper crust.
The five primitive dacites (>63 wt % SiO2) are interesting for various reasons, but we did not use them as parental liquids for most of the crystal fractionation and pyroxenite addition calculations because we do not believe primitive andesites and dacites can evolve by crystal fractionation to low Mg-number basalts and basaltic andesites. Two volcanic samples from Stuck Mountain are high Mg-number dacites (Mg-number 59·0–62·0; 67·5–68·0 wt % SiO2). One of these high Mg-number dacites [in addition to a dacite breccia (68·5 wt % SiO2, Mg-number 53·6) and a volcaniclastic sample from Plafker et al. (1989) from the same area] have trace-element chemistry with distinctly high La/Yb ratios that are different from those of the rest of the arc samples (Fig. 6a). These compositions imply either very different crystal fractionation processes (probably involving garnet) or a distinctly different parental melt, which in turn may have been derived from a source with residual garnet, or a source with different trace-element composition. Other samples with distinctive HREE concentrations are preserved in the Klanelneechena Klippe, where garnet- and pyroxene-bearing quartz diorites and tonalites are thrust over accretionary complex assemblages south of the Border Ranges Fault (Fig. 1). These plutonic rocks have distinctly HREE-enriched trace-element patterns and could represent cumulates or restites in equilibrium with LREE-enriched, HREE-depleted lavas (Kelemen et al., 2003a). Evolved, strongly HREE-depleted silicic melts could mix with primitive basalts to produce moderately HREE-depleted, high Mg-number andesites and dacites. The HREE-depleted volcanic samples are fairly rare and were not incorporated into our modeling of primary magma compositions or crystal fractionation trends for the bulk of the Talkeetna Arc samples.
Effects of crustal fractionation
Numerous petrological studies predict the existence of large sections of lower crustal cumulates to explain the observed geochemical variation in evolved volcanic rocks (e.g. Kay & Kay, 1985). Volcanic differentiation trends in several active arcs have been related through fractionation of lower crustal cumulates using the composition of xenoliths (e.g. Conrad & Kay, 1984). Exposures of complementary volcanic and plutonic sections of island arc crust from the Talkeetna Arc provide an exceptional opportunity to quantitatively test models relating volcanic suites and lower crustal cumulates through fractional crystallization. Burns (1985) recognized the similarities between the exposed layered gabbronorite from the Tazlina–Barnette area and plutonic xenoliths from island arc lavas, and cited this evidence in support of the plagioclase–orthopyroxene/olivine–augite–magnetite (POAM) fractionation model emphasized by Gill (1981). Gill (1981) proposed a dominant role for POAM fractionation to account for the derivation of the basalt–andesite–dacite–rhyolite series in island arc settings. The calculated phase proportions of minerals in our layered gabbronorite samples indicate that plagioclase + orthopyroxene + clinopyroxene + Fe–Ti oxide ± amphibole represent the bulk of the residual cumulates (Table 8).
Trace-element ratios allow a clear distinction between cumulate gabbroic rocks on the one hand, and evolved plutonic and volcanic samples on the other hand, and illustrate the effect of POAM fractionation in the Talkeetna Arc. The group of cumulate gabbronorites is easily distinguishable in Fig. 11 because it has lower silica contents than evolved plutonic and volcanic samples. The volume of cumulates in lower crustal exposures may be large enough to drive primitive basaltic liquid compositions to higher SiO2, as explored quantitatively in the least-squares fractional crystallization modeling section, below. With the exception of two high Sr/Y dacites (same as high La/Yb samples described above), trends in Sr/Y appear generally to be the result of plagioclase fractionation (Fig. 11). The variation in Ti, Zr and V within the suite of arc rocks is clearly related to the nature and proportion of crystallizing phases (Pearce & Norry, 1979). Crystallization of V-rich, Fe–Ti oxides within the layered gabbronorite is reflected by a trend of decreasing Ti/Zr and V/Ti for residual melts (e.g. Pearce & Norry, 1979; Nielsen et al., 1994). Zr concentrations appear to be controlled almost exclusively by fractionation (also see Fig. 6). Zr enrichment relative to Sm is often attributed to amphibole fractionation (Thirlwall et al., 1994), and Zr/Sm in the layered gabbronorites is distinctly lower than the other arc rocks. However, augite fractionation (Thirlwall et al., 1994) and magnetite fractionation (Tribuzio et al., 1999) may also exert an influence on the Zr/Sm ratio.
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Studies of inclusions from the Lesser Antilles document cumulate-textured xenoliths containing 5–15 modal % magnetite and attribute many of the chemical trends within suites of island arc rocks to magnetite fractionation (Arculus & Wills, 1980
). Magnetite–ilmenite gabbronorite is prevalent in lower crustal gabbronorites between Tazlina Lake and Barnette Creek. Whole-rock trace-element patterns for these cumulate rocks are enriched in Ti relative to the HREE, whereas volcanic and upper-level plutonic rocks, presumably the derivative liquids, are distinctly depleted in Ti (Fig. 6).
The precipitation of amphibole, in addition to POAM fractionation, is also often cited to explain geochemical variations within arc volcanic suites [e.g. Gill (1981) and references therein]. However, the pargasitic hornblende in lower crustal gabbronorite in the Talkeetna Arc appears to be primarily a late magmatic phase from progressive fluid enrichment (DeBari & Coleman, 1989), the result of subsolidus re-equilibration (Burns, 1985), from the reaction of clinopyroxene and melt as suggested by Foden & Green (1992), and/or formed by reactions between cumulus minerals and evolved melts ± aqueous fluids (Costa et al., 2002).
Plagioclase accumulation is an important process within island arc crust (e.g. Woodhead, 1988). Eu anomalies in island arc magmas are at least partly due to the accumulation of plagioclase (Vukadinovic, 1993). In the Talkeetna Arc, correlation between Eu/Eu* (log scale) and Yb(N) reflects the strong effect of plagioclase fractionation and minimal effect of trapped liquids on Yb concentrations (Fig. 12a). Eu and Sr anomalies show a clear positive correlation for the range of the Talkeetna Arc lithologies (Fig. 12b), indicating that plagioclase accumulation is largely responsible for enriched Sr concentrations in cumulate gabbronorites. However, primary Talkeetna magmas probably did have an initial enrichment in Sr/Nd compared with MORB (Kelemen et al., 2003a).
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[Pr(N) x Nd(N)] and Eu/Eu* = Eu(N)/Least-squares fractional crystallization modeling
Least-squares calculations are commonly used to quantify the possible relationship of a series of volcanic rocks through fractional crystallization (e.g. Arculus & Wills, 1980
). Cumulate sections from the middle and lower crust of the Talkeetna Arc provide a broad range of mineral compositions and proportions with which to quantify the possible relationship of volcanic and chilled mafic rock compositions through fractionation of observed cumulate gabbronorite. The goal of this modeling is to determine if the most primitive volcanic and chilled mafic rocks (Mg-number >60) can be related to more evolved volcanic rocks through fractional crystallization of gabbronorites, and to constrain the relative proportions of crystal and liquid products required by this hypothesis.
Initially, it is important to assess Fe/Mg and REE equilibrium between the cumulus phases and volcanic rocks to determine appropriate fractionation assemblages for the least-squares calculations. Using an Fe/Mg Kdcpx/liquid of 0·23, calculated with total Feliquid (Sisson & Grove, 1993), clinopyroxene Mg-number is plotted in Fig. 13, in a panel adjacent to a plot of Mg-number vs SiO2 for volcanic and chilled mafic rocks. Clinopyroxene from Talkeetna pyroxenite, plagioclase pyroxenite, garnet gabbro, basal gabbronorite, lower crustal gabbronorite, and mid-crustal amphibole gabbronorite covers most of the range of equilibrium compositions required for the lavas and mafic inclusions (Fig. 13). This diagram was used to evaluate potential equilibrium pairs of clinopyroxene-bearing rocks and corresponding liquid compositions. When available, REE clinopyroxene/liquid equilibria were assessed in conjunction with Fe/Mg equilibria. For example, in step 3 of the least-squares modeling, calculated liquid REE concentrations from clinopyroxene in lower crustal gabbronorite [0718A4; Dy(N) 17·7, Yb(N) 15·8] are close to REE abundances in the parent basaltic andesite [0709P2A; Dy(N) 18·3, Yb(N) 14·5].
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The compositional variation for a series of volcanic and chilled mafic rocks was modeled by relating observed compositions through fractionation of phases in cumulate gabbronorite. Parent/daughter pairs of liquid compositions were selected based on decreasing Mg-number, increasing SiO2, and increasing REE concentrations. Phase compositions (but not phase proportions) were chosen from cumulate Talkeetna samples with clinopyroxene in Fe/Mg exchange equilibrium (and, sometimes, REE exchange equilibrium) with the parent liquid. The outputs include per cent crystallized, weight proportions of fractionated phases, and residuals for each of the oxides. Solutions from three steps of least-squares fractionation modeling are shown in Table 9 and illustrated in Fig. 13b. Phase proportions from cumulate gabbronorites (from Table 8) are compared with model solutions from steps 2 and 3 (Fig. 13c). Observed and calculated REE provide an additional test for the three steps of the fractionation modeling (Fig. 13d).
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Differentiation of three consecutive parent/daughter pairs was modeled as follows.
Step 1. A primitive chilled mafic rock (1728A3) was selected as the parent composition for the first step based on its high Mg-number (61·9), despite its slightly LREE-depleted whole-rock REE pattern. Another primitive, chilled mafic rock (Mg-number 64·0) did not yield acceptable solutions for the first step of the series. Possible solutions include assemblages of cpx + opx + plag + sp (Table 9, with mineral abbreviations defined) and oliv + plag (not shown). Both yield solutions with 20% crystallization. The ol + plag solution requires proportions of 25% olivine and 75% plagioclase. Such troctolite compositions are not present in the Talkeetna Arc section. The cpx + opx + plag + sp solution includes a lot of aluminous spinel (18%), and only 32% plagioclase. However, plagioclase is necessary in this step to attain a solution (cpx + opx + sp or ol + cpx + opx + plag + sp are not workable assemblages). Even though phase proportions in the cpx + opx + plag + sp solution from this step represent a pyroxene-rich gabbro, the only rock with clinopyroxene Mg-number appropriate for Fe/Mg equilibrium with the parent composition in this step is a plagioclase pyroxenite. Plagioclase pyroxenite is rare in the Talkeetna Arc, but was sampled from Scarp Mountain. Sample 1709P11 is a spinel-rich, plagioclase pyroxenite with clinopyroxene Mg-numbers of 84·4–87·0 (Table 2), which partially bridge the gap in Mg-number between more common plagioclase-free pyroxenites (Mg-numbers 87–91) and basal gabbronorites (Mg-numbers 65–83) (Figs 2 and 13). We reiterate, however, that plagioclase pyroxenites are very rarely exposed in the Talkeetna Arc.
Step 2. The daughter composition from step 1 was used as the parent for step 2. The daughter liquid in step 2 was a basaltic andesite (0709P2A), chosen on the basis of lower Mg-number (49·4) and higher SiO2 and REE. Again, only clinopyroxene from the plagioclase pyroxenite has Mg-number high enough for Fe/Mg equilibrium with the parent composition in step 2. However, solutions using this clinopyroxene composition yield phase proportions that are consistent with calculated modes of cumulate gabbronorites and require 39% crystallization (Fig. 13c).
Step 3. The daughter composition from step 2 was used as the parent for step 3. The daughter liquid in step 3 was an andesite (0709P2C), again chosen on the basis of lower Mg-number (43·1) and higher SiO2 and REE. Using a variety of phase compositions observed in lower crustal gabbronorites, all chosen to be close to Fe/Mg and REE equilibrium with the parental liquid for step 3, yields model results that span the range of observed phase proportions in lower crustal gabbronorites. Solutions consistently yield extremely low residuals (
R2 = 0·0000–0·0015). Using different lower crustal gabbronorite phase compositions for the fractionation assemblage produces a range of outputs requiring 30–45% crystallization. However, phase assemblages yielding solutions at the lower end of this range (30% crystallization) have low proportions of orthopyroxene (<3 wt %), and slightly higher proportions of clinopyroxene (20–26 wt %) and hornblende (35–45 wt %) than observed gabbronorites. Solutions that require close to 45% crystallization (Table 9) have phase proportions that closely match observed phase proportions (Fig. 13; Table 8).
REE abundances for the three daughter compositions increase in each step. The calculated and observed REE patterns are similar for steps 2 and 3 (Fig. 13d). The calculated REE concentrations are dependent on the selection of crystal/liquid partition coefficients and phase proportions from solutions (which vary slightly in alternative solutions).
Cumulative results for the overall per cent crystallized and melt remaining for the three least-squares steps are also shown in Table 9 (bottom right of each step), assuming no eruption of the remaining liquid at each step (all remaining liquid is used for the parent magma in the subsequent step). These calculations indicate 73·1% crystallization (26·9% melt remaining), by weight, after the three steps.
Attempts to model the formation of more evolved andesites and dacites through fractional crystallization were unsuccessful. Solutions capable of producing dacites and rhyolites, using other evolved basaltic andesite and andesite compositions and a range of cumulate gabbro fractionation assemblages (e.g. opx + cpx + plag + amph + sp), were also unattainable. In addition, fractionation of observed phase compositions from high Mg-number andesites and dacites to form evolved andesite–dacite–rhyolite was surprisingly difficult to model. Using a high Mg-number dacite parent yielded one possible solution out of 10 parent–daughter combinations. However the solution, requiring 15% crystallization of the assemblage cpx + opx + plag + sp + hbl, would have minimal effect on La/Yb and is inconsistent with the difference between parent and daughter REE patterns.
The free energy minimization program MELTS (Ghiorso & Sack, 1995) was used as an independent method for comparison with results from least-squares fractional crystallization modeling. Liquid compositions from four fractional crystallization MELTS calculations at 0·5 and 1 GPa and temperatures from 1250 to 1000 °C are shown in Fig. 13b. Using the parent composition from step 1 of the least-squares fractionation modeling, MELTS models are shown for 2 wt % H2O and fO2 of NNO + 1 (where NNO is nickel–nickel oxide buffer) at 0·5 GPa, and for 1 wt % H2O and fO2 of FMQ + 1 (where FMQ is fayalite–magnetite–quartz buffer), 2 wt % H2O and FMQ + 2, and 3 wt % H2O and FMQ + 2 at 1 GPa. Although we tried other runs, the MELTS calculations presented in this figure have the closest results to either observed volcanic rock compositions or observed phases in the lower crustal gabbronorites for reasonable H2O contents and fO2.
The liquid compositions for runs with 2 and 3 wt % H2O at FMQ + 2 at 1 GPa are closest to the general trend of the volcanic rock compositions. However, fractionated solid assemblages from these two MELTS runs are not the same as observed phase proportions in the lower crustal gabbronorites or as results from least-squares fractional crystallization modeling. At 1 wt % H2O, FMQ + 1 and 2 wt % H2O, FMQ + 2, solid assemblages include sp + garnet + plag + cpx, whereas solid assemblages for 3 wt % H2O, FMQ + 2 include only sp + cpx.
For calculations with pressure set at 0·5 GPa, the fractionated solid proportions predicted in MELTS calculations are close to observed phase proportions in the lower crustal gabbronorites and to results from least-squares fractional crystallization modeling, but the evolving liquid composition from these MELTS runs does not match the observed trend of volcanic rock compositions as closely. Results for 2 wt % H2O, NNO + 1 and NNO + 2 at 0·5 GPa include fractionated solids with opx + cpx + plag + sp in proportions in the range observed in our samples.
Thus, using a variety of pressures, water contents and oxygen fugacities, we find that the MELTS calculations do not reproduce the liquid compositions and fractionated phase proportions observed in the Talkeetna section. In contrast, the least-squares models—by design—simultaneously fit both solid and liquid compositions better than the MELTS calculations. One possible conclusion is that the phase proportions produced by the least-squares models are not, in fact, equilibrium proportions of co-crystallizing minerals in arc magmas. However, given that isotope and petrological data indicate a genetic link between the plutonic and gabbroic rocks in the Talkeetna section, we believe that the least-squares models probably do yield realistic phase proportions, and thus provide better constraints on igneous differentiation within the Talkeetna section, compared with the MELTS calculations. This is understandable, given the very limited number of experimental data on water-undersaturated crystallization of primitive, basaltic magmas in the pressure range from 0·2 to 1·0 GPa that are available to calibrate thermodynamic parameters used in MELTS.
Pyroxenite fractionation
Even the most primitive chilled mafic rocks in the Talkeetna Arc are not in Fe/Mg equilibrium with residual mantle peridotite, and so are probably not primary, mantle-derived melts. Instead, they probably formed via fractionation of primitive cumulates from a primary magma. The obvious primitive cumulates in the Talkeetna Arc section are pyroxenites present between residual mantle peridotites and basal gabbronorites in the Tonsina area. Alternatively, DeBari & Sleep (1991) implied that primary Talkeetna magmas might have undergone extensive fractionation of olivine to form cumulate dunites. However, the proportion of dunites in the Talkeetna mantle section is small, and many of the dunites that are present may be replacive or the result of incomplete peritectic reactions, rather than cumulates. Primitive hydrous melts in equilibrium with mantle peridotite at 1 GPa are in a reaction relationship with olivine (peritectic reaction: olivine + liquid 
pyroxene) and are pyroxene saturated (e.g. Müntener et al., 2001). In keeping with this result, most of the websterites in the Tonsina area are olivine free. Further, Talkeetna mantle dunites do not have Mg-numbers <89, as required to explain the evolution of primary melts to the lower Mg-numbers observed in the most primitive Talkeetna lavas. Instead, Tonsina pyroxenites have pyroxene Mg-numbers that overlap those in the residual mantle peridotites. Specifically, pyroxenites in the Tonsina area of the Talkeetna Arc have clinopyroxene compositions with Mg-number 87–91 and 0·6–1·0 wt % Cr2O3 (DeBari & Coleman, 1989; R. Workman, unpublished data, 2003). Clinopyroxene in dunite and harzburgite has Mg-number 90–94, and slightly lower Cr2O3 contents (0·5–0·7 wt %) (DeBari & Coleman, 1989). Thus, we conclude that pyroxenites, not dunites, were the most primitive cumulates in the Talkeetna Arc section.
To account for the crystallization of pyroxene at the base of the crust and estimate a primary magma composition for the Talkeetna crustal section, we modeled pyroxenite fractionation by incrementally adding pyroxene and spinel to the most primitive liquid compositions obtained from lavas and mafic inclusions, until the estimated melt was in equilibrium with clinopyroxene with Mg-number of 92. The clinopyroxene Mg-number of 92 was chosen as a typical value for clinopyroxene in residual mantle peridotite. The most primitive chilled mafic rocks and lavas in the Talkeetna Arc are in equilibrium with clinopyroxene compositions with Mg-numbers of 85–88.
Although clinopyroxene-bearing orthopyroxenites are present in 10% of the outcrop area, the majority of websterites in the Tonsina area have 70–75% clinopyroxene, 25–30% orthopyroxene, and <2% spinel (Kelemen et al., 2003a; DeBari & Coleman, 1989). Websterite with these mineral proportions, in Fe/Mg exchange equilibrium with the initial melt, was added to two primitive chilled mafic rock and two basaltic lava compositions in increments of 0·2 wt %, while calculating a new clinopyroxene and orthopyroxene based on Fe/Mg equilibrium with the new melt composition at each step (using a constant Cr–Al spinel composition). Because of the uncertainties involving cpx Fe/Mg Kd values and Fe3+/Fe2+, two sets of values for each are used to show the range of possible results [cpx Fe2+/Mg Kd values of 0·27 (Sisson & Grove, 1993) and 0·33 (Müntener et al., 2001); Fe3+liquid of 0·14 and 0·20]. Estimated primary magma compositions and the weight per cent added pyroxenite are presented in Table 10.
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To attain a primary liquid composition in equilibrium with clinopyroxene having a Mg-number of 92, the required proportion of added pyroxenite is substantial. With a cpx Fe2+/Mg Kd of 0·27 and Fe3+liquid of 0·14, the proportion of pyroxenite (with phase proportions of 70% cpx:28% opx:2% sp) added to the four primitive compositions ranges from
26 to 33 wt % (Table 10). A slightly higher proportion of pyroxenite is required using 75% cpx:23% opx:2% sp for all the calculations. Varying the cpx/melt Kd from 0·27 to 0·33 has a dramatic effect, increasing the required proportion of pyroxenite by 15–18 wt %. With increasing Fe3+liquid (from 0·14 to 0·20), the required proportion of pyroxenite decreases by 5–8 wt %. If we stop at an estimated melt in equilibrium with clinopyroxene with Mg-number of 90, this decreases the required proportion of pyroxenite by 10 wt %. Using these different values is intended to provide bounds on the potential amount of pyroxenite fractionation. In almost every case where preferred values are used, >25 wt % pyroxenite is required to achieve a liquid composition in Fe/Mg exchange equilibrium with clinopyroxene in mantle peridotite (Table 10). Despite the substantial uncertainties involved in these calculations, this provides an approximate minimum estimate for the amount of pyroxenite fractionation that potentially occurred at the base of the arc crust.
The estimated primary magma compositions after pyroxenite addition calculations (Table 10) are similar to the calculated primary magma composition for the Talkeetna Arc of DeBari & Sleep (1991). Our calculated primary melts have 14–16 wt % Al2O3 and 11–12 wt % MgO. CaO contents range from 11 to 14 wt %, and are slightly higher than expected (13–14 wt %) for calculations using sample 1728A3 (Table 10). This sample, with 12 wt % CaO, probably has a small amount of accumulated plagioclase and/or clinopyroxene, which is also reflected in its whole-rock REE pattern. The calculated primary magma composition of DeBari & Sleep (1991) was a high-Mg (11·3 wt %), low-Al (15·0 wt %) basaltic composition (51·1 wt % SiO2) with a flat REE pattern (also listed in Table 10).
This is a dramatic corroboration of DeBari & Sleep's first-order approach. It should be recalled that we have derived our primary melt compositions from the most primitive observed lavas and mafic inclusions, considering only pyroxenite addition to reach a primary melt in Fe/Mg equilibrium with residual mantle peridotite. In contrast, DeBari & Sleep (1991) used a mass balance of all the exposed arc rocks, in their estimated proportions, to calculate an approximate bulk composition for the Talkeetna Arc crust. As such, they interpreted the bulk composition as the integrated mass flux from mantle to crust, after assuming that there was some dunite fractionation in the mantle. The trace-element and isotopic data presented in this paper are consistent with crystallization of the entire arc section along a single liquid line of descent from a homogeneous primary melt. Therefore, the bulk crust composition of DeBari & Sleep might be expected to correspond to the primary magma composition inferred from our data.
However, the apparent agreement between our estimates and those of DeBari & Sleep mask some essential differences. Whereas DeBari & Sleep assumed that the arc process produced 18 wt % ultramafic rocks (45·5 wt % dunite) below the Moho, with the remainder of the arc magma flux present as gabbroic and volcanic rocks in the crust, we infer that primary melts crystallized at least one-quarter of their mass as primitive pyroxenite cumulates, probably along the Moho, where appropriate pyroxenite compositions are exposed in the Talkeetna section. In fact, in comparison with DeBari & Sleep, we derive a very different composition for the gabbroic to volcanic crust overlying this Moho-level pyroxenite.
Estimates of the type and proportion of ultramafic cumulates in island arcs have important consequences. The proportion of ultramafic cumulates directly affects estimates of the bulk composition of arc crust. The degree of fractional crystallization from primary, mantle-derived melts below the Moho is also crucial to our understanding of the formation of high-Al basalts parental to plagioclase-bearing crustal rocks. Our calculations suggest that >25 wt % of primary melts crystallized as pyroxenites at the base of the crust. The experiments of Müntener et al. (2001) show that up to 50% of primary, mantle-derived melts could crystallize as pyroxenites prior to plagioclase saturation. If olivine is incorporated into our pyroxenite addition calculations, so ultramafic cumulate with between 10 and 80% olivine is added to the primitive liquid compositions, 18–30 wt % ultramafic cumulate is required to achieve a liquid composition in Fe/Mg exchange equilibrium with clinopyroxene in mantle peridotite. Previous studies using least-squares fractionation models require 21% (Conrad & Kay, 1984) and 16–26% (Gust & Perfit, 1987) crystallization of ultramafic cumulates to produce high-Al basalt in island arcs.
However, exposures of pyroxenites in the Tonsina area represent a thickness of <1 km. Thermobarometry on overlying garnet granulites requires that the crust above the pyroxenites was 30–35 km thick at one time (DeBari & Coleman, 1989). There are small ultramafic intrusions into the lower crustal gabbronorites at several places in the Talkeetna section (e.g. Rose, 1966; Clarke, 1972; Pavlis, 1983; Burns, 1985) but altogether pyroxenites represent <3% of the outcrop area in the plutonic parts of the Talkeetna Arc section. The discrepancy between the amount of pyroxenite fractionation deduced from crystallization modeling and the amount of pyroxenite observed in the Talkeetna section is probably not due to gaps in the section as a result of faults or missing outcrop. Although the total outcrop area is small, Moho-level outcrops in the Tonsina area expose several unfaulted, high-temperature contacts between underlying mantle peridotites, primitive pyroxenites (clinopyroxene Mg-number 90–85), and overlying gabbronorites (clinopyroxene Mg-number <85). Thus, the thin layer of pyroxenite is bounded on either side by high-temperature, arc-aged contacts with Talkeetna upper mantle and lowermost crust.
Instead, the discrepancy between the observed proportion of pyroxenites and the proportion required by crystal fractionation modeling may be explained by ‘delamination’ (Kay & Kay, 1985, 1988, 1990, 1991; Arndt & Goldstein, 1989; Turcotte, 1989; DeBari & Sleep, 1991; Jull & Kelemen, 2001). Although many treatments of ‘delamination’ have called upon the presence of abundant garnet in eclogites or garnet granulites, Arndt & Goldstein (1989) noted that ultramafic cumulates have lower Mg-number than residual mantle peridotites, and so are generally denser than underlying mantle. Jull & Kelemen (2001) and Müentener et al. (2001) quantified this. They showed that calculated densities for an olivine clinopyroxenite xenolith from the Aleutians (bulk Mg-number of 84) and an experimentally produced pyroxenite (bulk Mg-number of 82) are 50 kg/m3 higher than the density of fertile lherzolite, and 75 kg/m3 denser than depleted oceanic peridotite, at 1 GPa and temperatures from 900 to 1200 °C. These are the P–T conditions that were present beneath the Talkeetna Arc while it was magmatically active. At the higher end of this temperature range, the density contrast and relative viscosities would allow pyroxenites to delaminate almost as quickly as they formed (Jull & Kelemen, 2001).
Garnet granulites are found in a narrow horizon in the Talkeetna area, immediately above the pyroxenites (DeBari & Coleman, 1989). Jull & Kelemen (2001) found that garnet granulites with the composition of Tonsina garnet granulites would be 100 kg/m3 denser than the underlying mantle at Tonsina Moho P–T conditions. For a Moho temperature of 1000 °C, Jull & Kelemen (2001) found that a layer of ultramafic cumulates and garnet granulites varying in thickness from 1 to 2 km would become gravitationally unstable and ‘drip’ into the underlying mantle in a few thousand to a few million years (non-Newtonian dry olivine rheology with background strain rates from 10–14 to 10–18/s; abundant, dissolved H2O would shorten instability times still further).
Thus, the discrepancy between the observed proportion of pyroxenites (<5% of the arc section) and the proportion required by crystal fractionation modeling (>25%) is best understood as the result of gravitational instability, with dense ultramafic cumulates, probably together with dense garnet granulites, foundering into the underlying mantle during the time when the Talkeetna Arc was magmatically active, or in the initial phases of slow cooling (and sub-solidus garnet growth) immediately after the cessation of arc magmatism. Alternatively, either the small proportion of pyroxenite in Talkeetna Arc outcrops is not representative of the true proportion of rocks that originally formed the arc section, or pyroxenite ‘sills’ within the mantle wedge were present at depths greater than the maximum depth of exposure.
If the interpretation is correct that the bulk composition of island arc crust is close to the integrated mass flux from mantle to crust, then the bulk composition of the Talkeetna Arc crust is probably close to the composition of the most primitive arc lavas and mafic inclusions. Compared with the crustal bulk composition estimated by DeBari & Sleep (1991) (also shown in Table 6) the most primitive chilled mafic rocks have higher Mg-number and lower SiO2.
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CONCLUSION |
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The geochemistry of a diverse suite of volcanic and plutonic rocks from the Talkeetna Arc crust is consistent with the linked magmatic origin of these different lithologies. Exposures of layered, lower crustal gabbronorite represent cumulates that crystallized in equilibrium with liquids that were extracted and yielded volcanic rocks of the Talkeetna Arc. Fractional crystallization of layered gabbronorite in the middle and lower crust was the predominant process in the formation of the lithologically heterogeneous Talkeetna Arc crust. Plagioclase + orthopyroxene + clinopyroxene + Fe–Ti oxide ± amphibole lithologies represent the bulk of the cumulate rocks that are complementary to the erupted lavas and noncumulate plutonic rocks. The compositional variation of a series of volcanic and chilled mafic rocks can be modeled through fractionation of cumulus phases within the range of their observed proportions, with a few notable exceptions.
Primitive volcanic samples have 8 wt % MgO and REE concentrations similar to calculated parental magma compositions for the layered gabbronorites. However, magmas parental to layered gabbronorites had already been fractionated themselves, forming as a result of fractionation of pyroxenites from primary magmas at the base of the crust. Calculations indicate that more than 25 wt % of primary, mantle-derived melts in the Talkeetna Arc may have crystallized as pyroxenites below the Moho. The considerable discrepancy between these results and the observed proportion of pyroxenites in the Talkeetna section may be best explained by delamination of dense pyroxenites into the less dense, underlying residual mantle.
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SUPPLEMENTARY DATA |
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Supplementary data are available at Journal of Petrology online.
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ACKNOWLEDGEMENTS |
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This study was supported by National Science Foundation Grant EAR-9910899. Analytical work for this study was made possible with assistance from Karen Hanghøj, Nobu Shimizu, Graham Layne, Nilanjan Chatterjee, Scott Kuehner, Diane Johnson and Charles Knaack. Field work was assisted by Terry Pavlis, Greg Hirth, Brad Hacker, Luc Mehl, Matt Rioux, Jim Mattinson, and Nik Christensen. Assistance with back-fraction routines was provided by Bill Leeman. Additional electron microprobe data were furnished by Rhea Workman, Brad Hacker, and Luc Mehl. The authors appreciate the thorough and insightful reviews by Richard Arculus and Mike Dungan. A. Greene appreciates the encouragement and patience from James Scoates and Dominique Weis while revising the manuscript.
* Corresponding author. E-mail: agreene{at}eos.ubc.ca
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