Abstract

Geochemical studies of Plio-Quaternary volcanic rocks from the Valle de Bravo–Zitácuaro volcanic field (VBZ) in central Mexico indicate that slab melting plays a key role in the petrogenesis of the Trans-Mexican Volcanic Belt. Rocks from the VBZ are typical arc-related high-Mg andesites, but two different rock suites with distinct trace element patterns and isotopic compositions erupted concurrently in the area, with a trace element character that is also distinct from that of other Mexican volcanoes. The geochemical differences between the VBZ suites cannot be explained by simple crystal fractionation and/or crustal assimilation of a common primitive magma, but can be reconciled by the participation of different proportions of melts derived from the subducted basalt and sediments interacting with the mantle wedge. Sr/Y and Sr/Pb ratios of the VBZ rocks correlate inversely with Pb and Sr isotopic compositions, indicating that the Sr and Pb budgets are strongly controlled by melt additions from the subducted slab. In contrast, an inverse correlation between Pb(Th)/Nd and 143Nd/144Nd ratios, which extend to lower isotopic values than those for Pacific mid-ocean ridge basalts, indicates the participation of an enriched mantle wedge that is similar to the source of Mexican intraplate basalts. In addition, a systematic decrease in middle and heavy rare earth concentrations and Nb/Ta ratios with increasing SiO2 contents in the VBZ rocks is best explained if these elements are mobilized to some extent in the subduction flux, and suggests that slab partial fusion occurred under garnet amphibolite-facies conditions.

INTRODUCTION

Slab-derived fluxing plays an important role in the global geochemical cycle, and gives rise to the distinctive chemical compositions observed in arc magmas (Gill, 1981). However, the physical mechanism of element recycling at subduction zones remains controversial: is the subduction component a hydrous fluid, derived from the dehydration of the altered portions of the oceanic crust and its sedimentary cover? Or is it a water-rich silicate melt derived from the partial fusion of the subducted materials? Or, more complicated still, is it a combination of both hydrous fluids and silicate melts? Can we distinguish the chemical fingerprints of these two processes of element recycling in the geochemistry of arc magmas? And if a distinction can be drawn, then which reactions are involved, and how can we relate them to the thermal and tectonic conditions of a subduction zone? Arcs constructed over thick continental crusts provide additional complications because mantle-derived melts could stagnate, differentiate and assimilate pre-existing rocks at various depths along their journey to the surface, in a process that can obscure the chemical imprints of the subduction environment. Is it possible to distinguish the chemical contributions of the continental crust from those derived from the subduction flux? Answering all these questions is essential for understanding and quantifying the solid-earth geochemical cycle on a global scale, and they have implications for more complicated questions such as the origin of the continental crust.

In this contribution we provide new geochemical data from the Valle de Bravo–Zitácuaro volcanic field (VBZ), a Pliocene to Quaternary monogenetic field located in the central sector of the Trans-Mexican Volcanic Belt (TMVB). The rocks in this area have major element compositions that are typical of arc volcanic rocks worldwide, and are similar to andesitic rocks erupted from long-lived Mexican stratovolcanoes. They are thus representative of the current tectonic and geological conditions that generate andesitic magmas in the central sector of the TMVB, and are an excellent suite to address magma generation processes in a thick-crusted continental arc. We will show that crystal fractionation and/or crustal contamination did not play a significant role in their petrogenesis, but that the geochemical signatures are better explained by different proportions of slab-derived silicate melts reacting with the sub-arc mantle wedge.

TECTONIC AND GEOLOGICAL SETTING

The Mexican subduction zone is a natural laboratory to test the efficiency with which chemical elements are subducted and ultimately recycled into arc magmas because parameters that govern its thermal structure gradually change along-strike (Fig. 1). The age and convergence rate of the subducted Rivera and Cocos plates gradually increase eastward along the trench, as the dip angle decreases to a sub-horizontal plane (Pardo & Suárez, 1995). Although the Wadati–Benioff zone below central Mexico is poorly defined, it has been long considered that the current volcanic front emerges when the subducted slabs reach ∼100 km depth, despite significant differences in the subduction angle. Recent thermal models and new focal data compilations have suggested, however, that the slab depth below central Mexico could be as shallow as 70–80 km, and that slab melting could occur at those depths (Manea et al., 2004, 2005). Indeed, slab melting has recently been invoked to explain the origin of some Miocene and Quaternary rock sequences along the arc (Luhr, 2000; Gómez-Tuena et al., 2003; Martínez-Serrano et al., 2004).

Fig. 1.

Generalized map of the Trans-Mexican Volcanic Belt (dark shading) and main tectonic features of the Rivera and Cocos plates [modified from Pardo & Suárez (1995)]. Numbers separated by a comma along the trench indicate the age of the oceanic crust (in Ma) and the convergence rate (in cm/year), respectively. Contours represent the depth of the top of the subducted slab (segmented when inferred). The studied area is depicted by a box labelled Valle de Bravo–Zitácuaro volcanic field (VBZ). A representative section of the subducted sediments and altered oceanic crust was sampled by the Deep Sea Drilling Project at Site 487 (DSDP Site 487). Also shown are the locations of some important volcanic fields: Palma Sola volcanic field (Palma Sola), Los Tuxtlas volcanic field (Tuxtlas), Citlaltépetl volcano (Citlaltépetl), Popocatépetl volcano (Popocatépetl), Chichinautzin volcanic field (Chichinautzin), Michoacán–Guanajuato volcanic field (MGVF), Colima volcano (Colima), Mascota volcanic field (Mascota), and San Juan volcano (S. Juan).

Fig. 1.

Generalized map of the Trans-Mexican Volcanic Belt (dark shading) and main tectonic features of the Rivera and Cocos plates [modified from Pardo & Suárez (1995)]. Numbers separated by a comma along the trench indicate the age of the oceanic crust (in Ma) and the convergence rate (in cm/year), respectively. Contours represent the depth of the top of the subducted slab (segmented when inferred). The studied area is depicted by a box labelled Valle de Bravo–Zitácuaro volcanic field (VBZ). A representative section of the subducted sediments and altered oceanic crust was sampled by the Deep Sea Drilling Project at Site 487 (DSDP Site 487). Also shown are the locations of some important volcanic fields: Palma Sola volcanic field (Palma Sola), Los Tuxtlas volcanic field (Tuxtlas), Citlaltépetl volcano (Citlaltépetl), Popocatépetl volcano (Popocatépetl), Chichinautzin volcanic field (Chichinautzin), Michoacán–Guanajuato volcanic field (MGVF), Colima volcano (Colima), Mascota volcanic field (Mascota), and San Juan volcano (S. Juan).

The compositional variability and the mass of the sediment fill that is subducted are not well known, but evidence from gravimetric profiles indicates that the sediment column could be as thin as ∼17 m above the young Rivera plate, and that the sediment thickness gradually increases eastwards along the trench (Manea et al., 2003). In addition, Leg 66 of the Deep Sea Drilling Project (DSDP) retrieved six cores off the coast of Acapulco (Moore et al., 1982). DSDP Site 487 (location shown in Fig. 1) reached into the basaltic basement. The sedimentary column of Site 487 has been described and geochemically analyzed (Moore & Shipley, 1988; Plank & Langmuir, 1998; Verma, 2000b; LaGatta, 2003), and consists of two units: ∼100 m of continentally derived Quaternary hemipelagic sediments that rest above ∼70 m of late Miocene–Pliocene pelagic sediments. Fragments of altered oceanic crust (AOC) from Site 487 have been described and analyzed, and are mainly composed of an olivine- and plagioclase-bearing basalt with minor amounts of an aphyric basalt (Verma, 2000b; LaGatta, 2003).

The VBZ is currently one of the best studied areas of the TMVB. Recent publications have mapped and addressed the local geology, stratigraphy, petrography, and the major and trace element geochemistry of the volcanic successions (Capra et al., 1997; Blatter & Carmichael, 1998b; Blatter et al., 2001). Furthermore, primitive andesites from the VBZ have also been studied experimentally (Blatter & Carmichael, 2001). In summary, the volcanism in the area is characterized by a myriad of cinder cones, domes and isolated fissural lava flows of andesite–dacite compositions with ages that vary between 5 ka and 6 Ma. The experimentally determined hydrous phase equilibria of the VBZ andesites are suggestive of pre-eruptive H2O contents of 2·5–6 wt %, and the varied phenocryst assemblages in the lavas appear to be the result of water loss during decompression and ascent (Blatter & Carmichael, 2001). The excellent match between the amount of crystals in the natural rocks and the experiments strongly suggests that the lavas are true liquid compositions. Importantly, Blatter & Carmichael (2001) suggested a direct mantle origin for some of the VBZ andesites on the basis of the high Ni and Cr abundances of the volcanic rocks, the presence of nodules of amphibole-bearing peridotite in one of the andesitic lava flows (Blatter & Carmichael, 1998a), and the similarity of compositions between the VBZ lavas and experimental water-saturated lherzolite melts (Hirose, 1997).

SAMPLES AND ANALYTICAL METHODS

In this contribution we report comprehensive geochemical data for the youngest volcanic sequences of the Valle de Bravo Zitácuaro volcanic field (< 3·77 Ma), as geologically characterized by Capra et al. (1997) and Blatter et al. (2001). We deliberately exclude the older and more altered volcanic successions (6–3·77 Ma).

Major element compositions for most samples were obtained by X-ray fluorescence (XRF) using a Siemens SRS-3000 instrument at Instituto de Geología, UNAM, following procedures described elsewhere (Lozano et al., 1995). Trace element abundances were determined by inductively coupled plasma mass spectrometry (ICP-MS) using a VG-PQ2+ at the Lamont–Doherty Earth Observatory (LDEO), following sample preparation and measurement procedures described by Gómez-Tuena et al. (2003). Additional major and trace element data were obtained by ICP-MS at Activation Laboratories of Ancaster, Canada (http://www.actlabs.com/) employing lithium metaborate/tetraborate fusions. Boron concentrations on selected samples were obtained by prompt gamma neutron activation (PGNA) at Activation Laboratories. Reproducibility is better than 10% for Li, Be, Sc, Cr, Ni, Cu, Zn, Mo, Sb, Cs, Tl and Pb, and better than 3% for all other elements. Typical precision for B is 100% at 0·5 ppm, 15% at 5 ppm, and better than 5% at 50 ppm.

Sr, Nd and Pb isotopic compositions were determined by thermal ionization mass spectrometry (TIMS) at LDEO using a VG sector 54-30 system equipped with nine Faraday collectors, following sample preparation and measurement procedures described by Gómez-Tuena et al. (2003). 87Sr/ 86Sr ratios are normalized to 86Sr/ 88Sr = 0·1194 and adjusted to an NBS987 standard ratio of 0·710230. During two separated intervals of analyses, the measured values of the NBS987 standard were 87Sr/ 86Sr = 0·710245 ± 0·000016 (2σ, n = 4) and 0·710271 ± 0·000014 (2σ, n = 6). 143Nd/144Nd ratios were normalized to 146Nd/144Nd = 0·72190 and corrected to a La Jolla Standard value of 143Nd/144Nd = 0·511860. Over the course of this study, the measured 143Nd/144Nd ratio of La Jolla standard was 0·511836 ± 0·000013 (2σ, n = 15). Pb isotopic ratios were corrected for mass fractionation using the LDEO 207Pb–204Pb double spike. Using this technique, samples are analyzed twice: once mixed with the double spike and once without it. The fractionation corrected values were adjusted to NBS981 standard values of 206Pb/204Pb = 16·9356, 207Pb/204Pb = 15·4891, and 208Pb/204Pb = 36·7006 (Todt et al., 1996). Over the course of this study, the double-spike fractionation corrected Pb isotope ratios of the NBS-981 standard were 206Pb/204Pb = 16·9356 ± 0·0048 (143 ppm), 207Pb/204Pb = 15·4912 ± 0·0047 (152 ppm), and 208Pb/204Pb = 36·7025 ± 0·014 (191 ppm) (2σ, n = 13). Total procedural blanks between 300 and 500 pg for Pb were negligible compared with the concentrations of this element in the dissolved rock samples.

RESULTS

Detailed petrographic descriptions of the VBZ rock suite have been reported in several previous publications (Capra et al., 1997; Blatter & Carmichael, 1998a, 1998b; Blatter & Carmichael, 2001; Blatter et al., 2001), and the main phenocryst assemblage and modal proportions of the rocks analyzed in this study are given in Table 1. Major and trace element abundances and Sr–Nd–Pb isotopic compositions are reported in Tables 2 and 3. Overall, the VBZ rocks vary from highly porphyritic to nearly aphyric, with the least evolved andesites having a phenocryst mineralogy of olivine (Ol) and orthopyroxene (Opx), or clinopyroxene (Cpx) and Opx. More evolved andesites and dacites usually contain Opx, Cpx, occasional hornblende (Hbl) and rare biotite (Bi). Importantly, most of the lavas do not contain plagioclase as a phenocryst, a remarkable feature that was previously noted by Blatter & Carmichael (1998b). Nevertheless, plagioclase is commonly observed in the groundmass, forming a trachytic texture along with glass and opaque minerals. All kinds of rocks often have small but ubiquitous quartz xenocrysts; some quartzites and schists, which were apparently entrained from the locally exposed basement, are also included as xenoliths. These nodules are often surrounded by a pyroxene shield and/or a glassy aureole, but sharp edges along the contact with the host magma are also commonly observed. These features indicate that the xenoliths were rapidly incorporated into the ascending and cooling magma shortly before eruption.

Like many arcs around the world, the igneous rocks from the VBZ belong to the subalkaline magmatic series (Fig. 2a) and do not display FeO-total enrichment during differentiation, plotting within the calc-alkaline field in an AFM diagram (not shown). The major element compositions of the VBZ rocks are very similar to those observed in the large Mexican stratovolcanoes, classified mostly as andesites and dacites, but some of the least evolved rocks from the VBZ tend to have higher K2O contents at relatively low SiO2 contents (Fig. 2b). These latter lavas were classified as shoshonites by Blatter et al. (2001).

Table 1:

Modal mineralogy of the Valle de Bravo–Zitácuaro volcanic rocks

Sample Ol Pl Cpx Opx Hbl Bi Op Qtz-xen Gms 
VBZ-LRE          
Zit-99-1A – 0·1 1·7 – –  – 94·0 
Zit-99-2 – 1·7 – –  – 96·2 
Zit-99-3 0·2 – 0·7 1·5 – –  1·5 95·9 
Zit-99-5 – – –  0·1 –  – 99·9 
Zit-99-17 – 11·5 2·8 2·4 – –  – 83·2 
Zit-99-18 – 4·2 –   0·1 92·5 
Zit-99-20 – – 0·5 0·9 – –  – 98·3 
Zit-99-23 – 0·2 0·4 0·7 0·6 – 0·7 0·2 96·9 
Zit-99-28 – 0·1 0·2 1·3 – –  0·2 98·1 
Zit-99-31 2·5 0·1 – – – –  6·4 90·8 
Zit-99-30 0·1 – 0·7 1·5 – –  – 97·5 
VBZ-HRE          
Zit-99-25 – 0·5 1·6 – –  – 96·7 
Zit-99-8 3·9 – 1·5 0·5 – –  – 94·0 
Zit-99-9 0·1 – – 1·7 – –  0·7 97·5 
Zit-99-10 – – 0·1 3·5 2·2 – 0·3 – 93·6 
Zit-99-12 – – 0·5 1·5 – –  – 98·0 
Zit-99-14 5·9 – – 0·2 – –  0·5 93·3 
Zit-99-25 – 0·5 1·6 – –  – 96·7 
Zit-99-26 – – 0·4 – – – 1·5 94·0 
Sample Ol Pl Cpx Opx Hbl Bi Op Qtz-xen Gms 
VBZ-LRE          
Zit-99-1A – 0·1 1·7 – –  – 94·0 
Zit-99-2 – 1·7 – –  – 96·2 
Zit-99-3 0·2 – 0·7 1·5 – –  1·5 95·9 
Zit-99-5 – – –  0·1 –  – 99·9 
Zit-99-17 – 11·5 2·8 2·4 – –  – 83·2 
Zit-99-18 – 4·2 –   0·1 92·5 
Zit-99-20 – – 0·5 0·9 – –  – 98·3 
Zit-99-23 – 0·2 0·4 0·7 0·6 – 0·7 0·2 96·9 
Zit-99-28 – 0·1 0·2 1·3 – –  0·2 98·1 
Zit-99-31 2·5 0·1 – – – –  6·4 90·8 
Zit-99-30 0·1 – 0·7 1·5 – –  – 97·5 
VBZ-HRE          
Zit-99-25 – 0·5 1·6 – –  – 96·7 
Zit-99-8 3·9 – 1·5 0·5 – –  – 94·0 
Zit-99-9 0·1 – – 1·7 – –  0·7 97·5 
Zit-99-10 – – 0·1 3·5 2·2 – 0·3 – 93·6 
Zit-99-12 – – 0·5 1·5 – –  – 98·0 
Zit-99-14 5·9 – – 0·2 – –  0·5 93·3 
Zit-99-25 – 0·5 1·6 – –  – 96·7 
Zit-99-26 – – 0·4 – – – 1·5 94·0 

Modal proportions of phenocrysts (>0·3 mm) from a minimum of 1200 points. Ol, olivine; Pl, plagioclase; Cpx, clinopyroxene; Opx, orthopyroxene; Hbl, hornblende; Bi, biotite; Qtz-xen, quartz xenocrysts; Op, opaque minerals; Gms, groundmass.

Table 2:

Major and trace element compositions of the Valle de Bravo–Zitácuaro volcanic rocks

Sample: ZIT-99-1a ZIT-99-1B* ZIT-99-2 ZIT-99-3 ZIT-99-5 ZIT-99-17* ZIT-99-18 ZIT-99-19* ZIT-99-20* ZIT-99-23* 
Unit: VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE 
Long. W: 100·20 100·20 100·22 100·26 100·28 100·15 100·16 100·20 100·34 100·39 
Lat. N: 19·22 19·22 19·19 19·18 19·10 19·40 19·40 19·45 19·43 19·32 
Locality: C. El Rosario C. El Rosario Colorines Sto. Tomás Plátanos Agua Zarca Sta. María Delicias Arroyo Lindero La Dieta Pueblo Nuevo El Álamo 
SiO2 62·1 63·9 62·0 64·7 62·6 61·2 63·2 60·5 62·2 64·2 
TiO2 0·7 0·6 0·6 0·6 0·7 0·7 0·6 0·6 0·6 0·5 
Al2O3 16·1 15·9 16·1 16·0 17·7 16·6 16·9 18·4 16·3 16·4 
Fe2O3T 5·4 4·4 5·3 4·2 4·8 5·5 4·2 4·5 4·6 4·0 
MnO 0·08 0·07 0·08 0·06 0·07 0·08 0·07 0·07 0·07 0·07 
MgO 4·9 3·6 5·1 3·4 2·5 3·0 2·8 2·0 3·9 3·1 
CaO 5·9 5·1 5·8 5·4 5·3 5·6 5·1 4·5 5·5 4·6 
Na23·7 4·2 3·7 3·9 4·2 3·5 4·0 3·5 3·8 4·3 
K21·5 1·8 1·8 1·8 2·2 1·9 1·7 1·6 2·1 2·3 
P2O5 0·1 0·2 0·1 0·1 0·2 0·1 0·1 0·1 0·1 0·2 
LOI 0·3 0·3 0·0 0·0 0·0 1·5 0·9 4·7 0·9 0·0 
Total 100·7 100·0 100·6 100·0 100·3 99·7 99·5 100·6 100·2 99·5 
Sc  12         
104 78 115 85 104 82 74 52 82 51 
Cr  100    62   92 66 
Co 19·7 13·7 19·0 14·0 11·7 11·2 12·0 6·5 13·6 9·6 
Ni 120 55 50 61  24 34  49 54 
Cu 20 19 22 36 12 16 18  27 16 
Zn 88 53 64 66 59 52 43 47 53 48 
Ga 18·4 17·7 19·1 18·6 21·4 15·9 20·0 15·8 17·2 16·6 
Li 9·0  8·6 8·0 14·1  16·6    
Be 1·1 1·0 1·4 1·1 1·5 1·3 1·5    
10·6 8·1 3·0 5·9 6·8 3·8 9·9   9·0 
Rb 21·3 26·6 28·8 21·8 32·5 44·6 28·4 28·1 27·3 30·4 
Sr 797 683 778 878 799 471 828 457 754 699 
14·8 12·6 15·7 10·8 13·6 15·7 12·6 11·5 10·5 10·0 
Zr 103 125 127 106 149 126 108 105 107 121 
Nb 4·33 3·52 4·01 2·90 4·35 4·24 4·31 3·41 2·88 3·13 
Mo 0·81  0·60 0·56 0·83  0·49    
Sn 0·33  0·94 0·63 0·86 1·33 0·71    
Sb 0·02  0·06 0·05 0·10  0·07   1·99 
Cs 0·43 0·64 0·93 0·44 1·26 0·90 2·18 1·93 1·06 1·19 
Ba 364 382 391 367 453 361 489 601 420 473 
La 13·17 12·86 13·18 10·46 14·73 12·48 14·64 11·24 11·68 14·65 
Ce 28·30 26·43 28·31 22·13 31·57 25·31 30·11 20·55 23·99 29·43 
Pr 3·64 3·44 3·66 2·91 4·05 3·61 3·87 3·07 3·32 3·71 
Nd 15·03 13·36 15·08 12·12 16·66 14·39 15·53 11·83 13·17 13·78 
Sm 3·19 2·78 3·24 2·62 3·56 3·20 3·17 2·50 2·69 2·52 
Eu 1·02 0·97 1·01 0·88 1·13 0·93 1·03 0·88 0·91 0·85 
Gd 3·01 2·66 3·02 2·37 3·12 3·09 2·76 2·38 2·46 2·24 
Tb 0·46 0·40 0·47 0·35 0·45 0·48 0·41 0·36 0·36 0·32 
Dy 2·60 2·24 2·71 2·00 2·42 2·73 2·21 2·02 1·96 1·76 
Ho 0·51 0·43 0·54 0·38 0·48 0·52 0·43 0·39 0·36 0·33 
Er 1·42 1·31 1·49 1·05 1·30 1·55 1·16 1·17 1·07 1·02 
Tm  0·19    0·24  0·17 0·15 0·15 
Yb 1·33 1·20 1·50 1·06 1·25 1·44 1·17 1·10 1·00 1·00 
Lu 0·208 0·180 0·230 0·162 0·190 0·206 0·179 0·158 0·142 0·140 
Hf 2·93 3·25 3·52 3·01 3·98 3·28 3·15 2·86 2·87 3·12 
Ta 0·273 0·233 0·282 0·210 0·291 0·340 0·337 0·283 0·198 0·222 
0·143  0·226 0·109 0·179  0·124    
Tl 0·056 0·079 0·207 0·134 0·194 0·218 0·209 0·284 0·171 0·238 
Pb 4·00  10·54 5·79 7·44 5·29 7·24 7·91 6·31 7·51 
Th 2·19 3·02 3·33 1·97 3·58 3·67 2·93 2·85 2·40 3·32 
0·476 0·831 0·980 0·682 0·971 1·123 1·017 1·117 0·793 0·982 
Sample: ZIT-99-1a ZIT-99-1B* ZIT-99-2 ZIT-99-3 ZIT-99-5 ZIT-99-17* ZIT-99-18 ZIT-99-19* ZIT-99-20* ZIT-99-23* 
Unit: VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE 
Long. W: 100·20 100·20 100·22 100·26 100·28 100·15 100·16 100·20 100·34 100·39 
Lat. N: 19·22 19·22 19·19 19·18 19·10 19·40 19·40 19·45 19·43 19·32 
Locality: C. El Rosario C. El Rosario Colorines Sto. Tomás Plátanos Agua Zarca Sta. María Delicias Arroyo Lindero La Dieta Pueblo Nuevo El Álamo 
SiO2 62·1 63·9 62·0 64·7 62·6 61·2 63·2 60·5 62·2 64·2 
TiO2 0·7 0·6 0·6 0·6 0·7 0·7 0·6 0·6 0·6 0·5 
Al2O3 16·1 15·9 16·1 16·0 17·7 16·6 16·9 18·4 16·3 16·4 
Fe2O3T 5·4 4·4 5·3 4·2 4·8 5·5 4·2 4·5 4·6 4·0 
MnO 0·08 0·07 0·08 0·06 0·07 0·08 0·07 0·07 0·07 0·07 
MgO 4·9 3·6 5·1 3·4 2·5 3·0 2·8 2·0 3·9 3·1 
CaO 5·9 5·1 5·8 5·4 5·3 5·6 5·1 4·5 5·5 4·6 
Na23·7 4·2 3·7 3·9 4·2 3·5 4·0 3·5 3·8 4·3 
K21·5 1·8 1·8 1·8 2·2 1·9 1·7 1·6 2·1 2·3 
P2O5 0·1 0·2 0·1 0·1 0·2 0·1 0·1 0·1 0·1 0·2 
LOI 0·3 0·3 0·0 0·0 0·0 1·5 0·9 4·7 0·9 0·0 
Total 100·7 100·0 100·6 100·0 100·3 99·7 99·5 100·6 100·2 99·5 
Sc  12         
104 78 115 85 104 82 74 52 82 51 
Cr  100    62   92 66 
Co 19·7 13·7 19·0 14·0 11·7 11·2 12·0 6·5 13·6 9·6 
Ni 120 55 50 61  24 34  49 54 
Cu 20 19 22 36 12 16 18  27 16 
Zn 88 53 64 66 59 52 43 47 53 48 
Ga 18·4 17·7 19·1 18·6 21·4 15·9 20·0 15·8 17·2 16·6 
Li 9·0  8·6 8·0 14·1  16·6    
Be 1·1 1·0 1·4 1·1 1·5 1·3 1·5    
10·6 8·1 3·0 5·9 6·8 3·8 9·9   9·0 
Rb 21·3 26·6 28·8 21·8 32·5 44·6 28·4 28·1 27·3 30·4 
Sr 797 683 778 878 799 471 828 457 754 699 
14·8 12·6 15·7 10·8 13·6 15·7 12·6 11·5 10·5 10·0 
Zr 103 125 127 106 149 126 108 105 107 121 
Nb 4·33 3·52 4·01 2·90 4·35 4·24 4·31 3·41 2·88 3·13 
Mo 0·81  0·60 0·56 0·83  0·49    
Sn 0·33  0·94 0·63 0·86 1·33 0·71    
Sb 0·02  0·06 0·05 0·10  0·07   1·99 
Cs 0·43 0·64 0·93 0·44 1·26 0·90 2·18 1·93 1·06 1·19 
Ba 364 382 391 367 453 361 489 601 420 473 
La 13·17 12·86 13·18 10·46 14·73 12·48 14·64 11·24 11·68 14·65 
Ce 28·30 26·43 28·31 22·13 31·57 25·31 30·11 20·55 23·99 29·43 
Pr 3·64 3·44 3·66 2·91 4·05 3·61 3·87 3·07 3·32 3·71 
Nd 15·03 13·36 15·08 12·12 16·66 14·39 15·53 11·83 13·17 13·78 
Sm 3·19 2·78 3·24 2·62 3·56 3·20 3·17 2·50 2·69 2·52 
Eu 1·02 0·97 1·01 0·88 1·13 0·93 1·03 0·88 0·91 0·85 
Gd 3·01 2·66 3·02 2·37 3·12 3·09 2·76 2·38 2·46 2·24 
Tb 0·46 0·40 0·47 0·35 0·45 0·48 0·41 0·36 0·36 0·32 
Dy 2·60 2·24 2·71 2·00 2·42 2·73 2·21 2·02 1·96 1·76 
Ho 0·51 0·43 0·54 0·38 0·48 0·52 0·43 0·39 0·36 0·33 
Er 1·42 1·31 1·49 1·05 1·30 1·55 1·16 1·17 1·07 1·02 
Tm  0·19    0·24  0·17 0·15 0·15 
Yb 1·33 1·20 1·50 1·06 1·25 1·44 1·17 1·10 1·00 1·00 
Lu 0·208 0·180 0·230 0·162 0·190 0·206 0·179 0·158 0·142 0·140 
Hf 2·93 3·25 3·52 3·01 3·98 3·28 3·15 2·86 2·87 3·12 
Ta 0·273 0·233 0·282 0·210 0·291 0·340 0·337 0·283 0·198 0·222 
0·143  0·226 0·109 0·179  0·124    
Tl 0·056 0·079 0·207 0·134 0·194 0·218 0·209 0·284 0·171 0·238 
Pb 4·00  10·54 5·79 7·44 5·29 7·24 7·91 6·31 7·51 
Th 2·19 3·02 3·33 1·97 3·58 3·67 2·93 2·85 2·40 3·32 
0·476 0·831 0·980 0·682 0·971 1·123 1·017 1·117 0·793 0·982 
Sample: ZIT-99-28* ZIT-99-29* ZIT-99-27* ZIT-99-31 ZIT-99-30 ZIT-99-25 ZIT-99-7* ZIT-99-8 ZIT-99-9 ZIT-99-10 
Unit: VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-HRE VBZ-HRE VBZ-HRE VBZ-HRE VBZ-HRE 
Long. W: 100·43 100·47 100·44 100·49 100·49 100·47 100·07 100·06 100·09 100·10 
Lat. N: 19·39 19·49 19·41 19·41 19·42 19·44 19·07 19·05 19·03 19·04 
Locality: C. Los Coyotes La Garita El Asoleadero C. El Molcajete La Soledad Las Pilas Sn. Fr. La Albarrada Temascaltepec C. Tezontle El Peñón 
SiO2 60·5 54·6 60·8 54·2 59·8 58·0 51·6 54·2 58·8 59·3 
TiO2 0·7 1·3 0·8 1·1 0·8 0·9 0·8 0·9 0·9 0·9 
Al2O3 16·3 16·9 16·0 17·0 16·3 15·1 16·3 15·9 17·2 16·1 
Fe2O3T 5·3 8·4 5·2 8·0 5·7 5·6 7·6 7·4 6·2 6·4 
MnO 0·09 0·12 0·09 0·13 0·09 0·08 0·13 0·11 0·09 0·10 
MgO 4·1 5·8 4·2 6·7 4·9 6·2 8·7 7·5 4·4 4·2 
CaO 6·1 7·4 6·3 7·1 5·6 7·8 8·8 7·8 6·1 6·8 
Na23·3 3·6 3·6 3·5 4·0 3·5 3·2 3·6 3·9 3·6 
K22·3 1·4 2·0 1·4 1·7 1·8 2·0 2·4 2·1 1·8 
P2O5 0·1 0·3 0·2 0·2 0·2 0·3 0·4 0·4 0·3 0·3 
LOI 0·8 −0·1 0·9 0·4 0·6 0·6 0·3 0·4 0·7 0·7 
Total 99·8 99·8 100·0 99·6 99·7 99·9 99·9 100·5 100·8 100·2 
Sc   17    24    
128 133 131 144 86 118 184 159 114 144 
Cr  129 106    381    
Co 16·3 26·1 16·3 28·3 19·9 24·3 32·2 31·5 18·4 20·9 
Ni 45 80 67 153 140 155 167 201 88 66 
Cu 23 31 25 34 30 42 46 48 31 47 
Zn 47 83 60 70 87 47 74 84 64 89 
Ga 18·9 17·7 18·0 18·4 18·9 20·1 15·7 18·9 20·3 20·7 
Li 7·1   10·4 10·2 9·5  13·0 13·4 9·0 
Be 1·2 1·3 1·0 1·3 1·5 1·6 1·0 1·9 2·2 1·8 
7·0 7·7 5·2  9·8 5·5 3·8 3·8 12·2 7·1 
Rb 23·5 21·7 22·9 21·0 30·7 24·6 35·3 37·6 31·3 25·5 
Sr 890 576 863 653 655 1847 818 1088 1002 1244 
15·8 22·2 15·1 23·2 15·1 16·7 18·6 23·1 29·9 28·9 
Zr 106 164 110 144 138 153 133 171 188 162 
Nb 3·04 7·35 3·19 10·87 4·89 4·21 3·14 5·90 7·08 4·67 
Mo 0·56   1·19 0·61 0·75  0·92 1·07 0·64 
Sn 0·62 1·20  0·90 0·94 0·95  1·04 1·15 1·15 
Sb 0·04   0·03 0·06 0·04 0·28 0·04 0·12 0·06 
Cs 0·91 0·57 0·79 0·60 0·87 0·70 1·38 0·80 1·16 0·89 
Ba 379 381 398 342 472 556 654 895 729 469 
La 11·87 16·27 11·58 14·91 14·17 33·15 30·86 44·87 55·14 35·32 
Ce 25·51 34·24 24·99 32·39 29·91 73·94 62·33 91·89 102·98 58·78 
Pr 3·58 4·71 3·45 4·27 4·01 9·86 8·37 12·04 14·28 9·49 
Nd 15·22 18·88 14·16 18·00 16·94 40·49 32·35 48·76 56·72 39·44 
Sm 3·38 4·18 3·08 4·06 3·58 7·03 6·04 8·92 10·17 7·14 
Eu 1·11 1·38 1·10 1·36 1·18 2·01 1·80 2·46 2·88 2·07 
Gd 3·11 4·10 2·95 4·01 3·24 5·18 4·99 6·82 8·16 6·34 
Tb 0·47 0·67 0·47 0·64 0·48 0·69 0·66 0·90 1·14 0·89 
Dy 2·85 3·82 2·72 3·85 2·70 3·14 3·45 4·24 5·63 4·91 
Ho 0·54 0·72 0·53 0·78 0·51 0·57 0·64 0·77 1·03 0·95 
Er 1·54 2·17 1·56 2·23 1·41 1·42 1·85 2·01 2·68 2·58 
Tm  0·32 0·24    0·26    
Yb 1·55 1·94 1·48 2·17 1·37 1·38 1·69 1·80 2·50 2·48 
Lu 0·242 0·285 0·222 0·338 0·203 0·202 0·242 0·271 0·375 0·366 
Hf 3·04 3·70 2·89 3·40 3·59 4·21 3·33 4·23 4·62 4·54 
Ta 0·206 0·512 0·169 0·626 0·332 0·255 0·183 0·326 0·421 0·316 
0·097   0·185 0·155 0·108  0·207 0·582 0·247 
Tl 0·131 0·133 0·128 0·109 0·237 0·141 0·164 0·225 0·216 0·088 
Pb 6·03 5·88 5·49 4·65 7·41 7·84 7·96 11·53 9·13 5·87 
Th 2·21 2·43 2·22 2·04 2·13 3·87 6·00 6·00 4·78 3·03 
0·772 0·724 0·751 0·592 0·867 1·140 1·838 1·942 1·542 1·015 
Sample: ZIT-99-28* ZIT-99-29* ZIT-99-27* ZIT-99-31 ZIT-99-30 ZIT-99-25 ZIT-99-7* ZIT-99-8 ZIT-99-9 ZIT-99-10 
Unit: VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-LRE VBZ-HRE VBZ-HRE VBZ-HRE VBZ-HRE VBZ-HRE 
Long. W: 100·43 100·47 100·44 100·49 100·49 100·47 100·07 100·06 100·09 100·10 
Lat. N: 19·39 19·49 19·41 19·41 19·42 19·44 19·07 19·05 19·03 19·04 
Locality: C. Los Coyotes La Garita El Asoleadero C. El Molcajete La Soledad Las Pilas Sn. Fr. La Albarrada Temascaltepec C. Tezontle El Peñón 
SiO2 60·5 54·6 60·8 54·2 59·8 58·0 51·6 54·2 58·8 59·3 
TiO2 0·7 1·3 0·8 1·1 0·8 0·9 0·8 0·9 0·9 0·9 
Al2O3 16·3 16·9 16·0 17·0 16·3 15·1 16·3 15·9 17·2 16·1 
Fe2O3T 5·3 8·4 5·2 8·0 5·7 5·6 7·6 7·4 6·2 6·4 
MnO 0·09 0·12 0·09 0·13 0·09 0·08 0·13 0·11 0·09 0·10 
MgO 4·1 5·8 4·2 6·7 4·9 6·2 8·7 7·5 4·4 4·2 
CaO 6·1 7·4 6·3 7·1 5·6 7·8 8·8 7·8 6·1 6·8 
Na23·3 3·6 3·6 3·5 4·0 3·5 3·2 3·6 3·9 3·6 
K22·3 1·4 2·0 1·4 1·7 1·8 2·0 2·4 2·1 1·8 
P2O5 0·1 0·3 0·2 0·2 0·2 0·3 0·4 0·4 0·3 0·3 
LOI 0·8 −0·1 0·9 0·4 0·6 0·6 0·3 0·4 0·7 0·7 
Total 99·8 99·8 100·0 99·6 99·7 99·9 99·9 100·5 100·8 100·2 
Sc   17    24    
128 133 131 144 86 118 184 159 114 144 
Cr  129 106    381    
Co 16·3 26·1 16·3 28·3 19·9 24·3 32·2 31·5 18·4 20·9 
Ni 45 80 67 153 140 155 167 201 88 66 
Cu 23 31 25 34 30 42 46 48 31 47 
Zn 47 83 60 70 87 47 74 84 64 89 
Ga 18·9 17·7 18·0 18·4 18·9 20·1 15·7 18·9 20·3 20·7 
Li 7·1   10·4 10·2 9·5  13·0 13·4 9·0 
Be 1·2 1·3 1·0 1·3 1·5 1·6 1·0 1·9 2·2 1·8 
7·0 7·7 5·2  9·8 5·5 3·8 3·8 12·2 7·1 
Rb 23·5 21·7 22·9 21·0 30·7 24·6 35·3 37·6 31·3 25·5 
Sr 890 576 863 653 655 1847 818 1088 1002 1244 
15·8 22·2 15·1 23·2 15·1 16·7 18·6 23·1 29·9 28·9 
Zr 106 164 110 144 138 153 133 171 188 162 
Nb 3·04 7·35 3·19 10·87 4·89 4·21 3·14 5·90 7·08 4·67 
Mo 0·56   1·19 0·61 0·75  0·92 1·07 0·64 
Sn 0·62 1·20  0·90 0·94 0·95  1·04 1·15 1·15 
Sb 0·04   0·03 0·06 0·04 0·28 0·04 0·12 0·06 
Cs 0·91 0·57 0·79 0·60 0·87 0·70 1·38 0·80 1·16 0·89 
Ba 379 381 398 342 472 556 654 895 729 469 
La 11·87 16·27 11·58 14·91 14·17 33·15 30·86 44·87 55·14 35·32 
Ce 25·51 34·24 24·99 32·39 29·91 73·94 62·33 91·89 102·98 58·78 
Pr 3·58 4·71 3·45 4·27 4·01 9·86 8·37 12·04 14·28 9·49 
Nd 15·22 18·88 14·16 18·00 16·94 40·49 32·35 48·76 56·72 39·44 
Sm 3·38 4·18 3·08 4·06 3·58 7·03 6·04 8·92 10·17 7·14 
Eu 1·11 1·38 1·10 1·36 1·18 2·01 1·80 2·46 2·88 2·07 
Gd 3·11 4·10 2·95 4·01 3·24 5·18 4·99 6·82 8·16 6·34 
Tb 0·47 0·67 0·47 0·64 0·48 0·69 0·66 0·90 1·14 0·89 
Dy 2·85 3·82 2·72 3·85 2·70 3·14 3·45 4·24 5·63 4·91 
Ho 0·54 0·72 0·53 0·78 0·51 0·57 0·64 0·77 1·03 0·95 
Er 1·54 2·17 1·56 2·23 1·41 1·42 1·85 2·01 2·68 2·58 
Tm  0·32 0·24    0·26    
Yb 1·55 1·94 1·48 2·17 1·37 1·38 1·69 1·80 2·50 2·48 
Lu 0·242 0·285 0·222 0·338 0·203 0·202 0·242 0·271 0·375 0·366 
Hf 3·04 3·70 2·89 3·40 3·59 4·21 3·33 4·23 4·62 4·54 
Ta 0·206 0·512 0·169 0·626 0·332 0·255 0·183 0·326 0·421 0·316 
0·097   0·185 0·155 0·108  0·207 0·582 0·247 
Tl 0·131 0·133 0·128 0·109 0·237 0·141 0·164 0·225 0·216 0·088 
Pb 6·03 5·88 5·49 4·65 7·41 7·84 7·96 11·53 9·13 5·87 
Th 2·21 2·43 2·22 2·04 2·13 3·87 6·00 6·00 4·78 3·03 
0·772 0·724 0·751 0·592 0·867 1·140 1·838 1·942 1·542 1·015 
Sample: ZIT-99-11* ZIT-99-12 ZIT-99-14 ZIT-99-24* ZIT-99-26      
Unit: VBZ-HRE VBZ-HRE VBZ-HRE VBZ-HRE VBZ-HRE      
Long. W: 100·14 100·19 100·25 100·45 100·46      
Lat. N: 19·03 19·05 19·07 19·46 19·43      
Locality: C. Pelón Potrero de Tenayac C. Pelón Agua Salada Las Pilas      
SiO2 52·6 56·4 54·9 55·1 55·7      
TiO2 1·0 0·9 0·8 1·3 1·1      
Al2O3 15·0 16·4 15·4 16·9 15·5      
Fe2O3T 7·4 6·6 7·4 7·9 7·0      
MnO 0·12 0·09 0·11 0·12 0·10      
MgO 8·9 5·1 8·9 4·8 5·6      
CaO 9·0 6·7 7·4 7·3 6·8      
Na23·6 3·9 3·2 3·7 3·6      
K21·5 2·5 1·3 1·4 2·8      
P2O5 0·4 0·4 0·2 0·4 0·4      
LOI 0·2 0·7 0·4 0·6 0·9      
Total 99·5 99·6 100·2 99·4 99·6      
Sc 22   19  
153 135 150 150 151      
Cr 438   106  
Co 33·6 22·2 33·6 23·2 32·2      
Ni 184 90 223 48 182      
Cu 43 43 47 28 34      
Zn 73 91 72 94 80      
Ga 16·6 21·0 17·2 19·6 20·5      
Li  11·8 6·2  9·7      
Be 1·1 2·7 1·3 1·2 2·8      
1·9 9·2 9·0 5·5 7·9      
Rb 20·3 37·8 21·2 25·4 52·8      
Sr 1230 1032 770 631 1301      
19·2 38·9 19·8 23·7 32·9      
Zr 121 177 137 199 246      
Nb 4·12 6·06 4·81 9·32 4·92      
Mo  0·98 0·76  0·79      
Sn 1·25 1·12 0·77 1·44 1·38      
Sb  0·18 0·10  0·05      
Cs 0·53 1·47 1·11 0·67 0·96      
Ba 590 887 434 425 993      
La 30·26 79·51 24·47 20·84 46·85      
Ce 63·25 155·45 47·88 43·75 95·78      
Pr 8·64 19·78 6·69 5·97 12·64      
Nd 33·09 79·58 26·88 23·75 51·90      
Sm 6·17 14·71 5·16 5·08 9·80      
Eu 1·82 3·82 1·56 1·62 2·79      
Gd 4·98 12·10 4·48 4·82 8·16      
Tb 0·66 1·61 0·64 0·74 1·13      
Dy 3·53 8·24 3·40 4·25 5·89      
Ho 0·63 1·44 0·67 0·80 1·06      
Er 1·84 3·64 1·78 2·43 2·75      
Tm 0·25   0·36  
Yb 1·61 3·26 1·69 2·20 2·54      
Lu 0·233 0·459 0·260 0·319 0·361      
Hf 3·09 4·80 3·38 4·42 6·61      
Ta 0·245 0·377 0·285 0·644 0·353      
 0·193 0·197  0·169      
Tl  0·271 0·238 0·201 0·280      
Pb 5·89 10·57 5·76 6·73 8·57      
Th 3·56 5·50 2·99 2·58 4·07      
1·059 1·779 0·943 0·783 1·704      
Sample: ZIT-99-11* ZIT-99-12 ZIT-99-14 ZIT-99-24* ZIT-99-26      
Unit: VBZ-HRE VBZ-HRE VBZ-HRE VBZ-HRE VBZ-HRE      
Long. W: 100·14 100·19 100·25 100·45 100·46      
Lat. N: 19·03 19·05 19·07 19·46 19·43      
Locality: C. Pelón Potrero de Tenayac C. Pelón Agua Salada Las Pilas      
SiO2 52·6 56·4 54·9 55·1 55·7      
TiO2 1·0 0·9 0·8 1·3 1·1      
Al2O3 15·0 16·4 15·4 16·9 15·5      
Fe2O3T 7·4 6·6 7·4 7·9 7·0      
MnO 0·12 0·09 0·11 0·12 0·10      
MgO 8·9 5·1 8·9 4·8 5·6      
CaO 9·0 6·7 7·4 7·3 6·8      
Na23·6 3·9 3·2 3·7 3·6      
K21·5 2·5 1·3 1·4 2·8      
P2O5 0·4 0·4 0·2 0·4 0·4      
LOI 0·2 0·7 0·4 0·6 0·9      
Total 99·5 99·6 100·2 99·4 99·6      
Sc 22   19  
153 135 150 150 151      
Cr 438   106  
Co 33·6 22·2 33·6 23·2 32·2      
Ni 184 90 223 48 182      
Cu 43 43 47 28 34      
Zn 73 91 72 94 80      
Ga 16·6 21·0 17·2 19·6 20·5      
Li  11·8 6·2  9·7      
Be 1·1 2·7 1·3 1·2 2·8      
1·9 9·2 9·0 5·5 7·9      
Rb 20·3 37·8 21·2 25·4 52·8      
Sr 1230 1032 770 631 1301      
19·2 38·9 19·8 23·7 32·9      
Zr 121 177 137 199 246      
Nb 4·12 6·06 4·81 9·32 4·92      
Mo  0·98 0·76  0·79      
Sn 1·25 1·12 0·77 1·44 1·38      
Sb  0·18 0·10  0·05      
Cs 0·53 1·47 1·11 0·67 0·96      
Ba 590 887 434 425 993      
La 30·26 79·51 24·47 20·84 46·85      
Ce 63·25 155·45 47·88 43·75 95·78      
Pr 8·64 19·78 6·69 5·97 12·64      
Nd 33·09 79·58 26·88 23·75 51·90      
Sm 6·17 14·71 5·16 5·08 9·80      
Eu 1·82 3·82 1·56 1·62 2·79      
Gd 4·98 12·10 4·48 4·82 8·16      
Tb 0·66 1·61 0·64 0·74 1·13      
Dy 3·53 8·24 3·40 4·25 5·89      
Ho 0·63 1·44 0·67 0·80 1·06      
Er 1·84 3·64 1·78 2·43 2·75      
Tm 0·25   0·36  
Yb 1·61 3·26 1·69 2·20 2·54      
Lu 0·233 0·459 0·260 0·319 0·361      
Hf 3·09 4·80 3·38 4·42 6·61      
Ta 0·245 0·377 0·285 0·644 0·353      
 0·193 0·197  0·169      
Tl  0·271 0·238 0·201 0·280      
Pb 5·89 10·57 5·76 6·73 8·57      
Th 3·56 5·50 2·99 2·58 4·07      
1·059 1·779 0·943 0·783 1·704      

Major elements (in wt %) were analyzed by XRF. Trace elements (in ppm) were analyzed by ICP-MS.

*Samples analyzed by ICP-MS for major and trace elements at Activation Laboratories, Canada. Boron analyzed by prompt gamma neutron activation (PGNA) at Activation Laboratories, Canada.

Table 3:

Sr, Nd and Pb isotope compositions of the Valle de Bravo–Zitácuaro volcanic rocks

 87Sr/86Sr 2σ 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb 143Nd/144Nd 2σ 
VBZ-LRE        
ZIT-99-1a 0·703377 12 18·5921 15·5654 38·2915   
ZIT-99-2 0·703469 10 18·6342 15·5775 38·3526 0·512832 16 
ZIT-99-3 0·703223 10 18·5605 15·5547 38·2151 0·512891 16 
ZIT-99-5 0·703523 10 18·5847 15·5609 38·2773 0·512833 20 
ZIT-99-28 0·703397 12 18·5598 15·5597 38·2240 0·512914 10 
ZIT-99-31 0·703663 10 18·6462 15·5776 38·3761 0·512853 12 
ZIT-99-30 0·704321 12 18·6372 15·5822 38·3509 0·512794 16 
VBZ-HRE        
ZIT-99-25 0·703306 10 18·5713 15·5584 38·2462 0·512978 
ZIT-99-8 0·704355 10 18·6755 15·5879 38·4227 0·512893 18 
ZIT-99-9 0·703847 10 18·6489 15·5787 38·3629 0·512891 12 
ZIT-99–10 0·703351 12 18·6019 15·5627 38·2872 0·512930 14 
ZIT-99-12 0·704253 10 18·6554 15·5782 38·3770 0·512876 14 
ZIT-99-14 0·703692 12 18·6362 15·5710 38·3424 0·512899 18 
ZIT-99-26 0·704217 10 18·6324 15·5759 38·3326 0·512876 12 
 87Sr/86Sr 2σ 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb 143Nd/144Nd 2σ 
VBZ-LRE        
ZIT-99-1a 0·703377 12 18·5921 15·5654 38·2915   
ZIT-99-2 0·703469 10 18·6342 15·5775 38·3526 0·512832 16 
ZIT-99-3 0·703223 10 18·5605 15·5547 38·2151 0·512891 16 
ZIT-99-5 0·703523 10 18·5847 15·5609 38·2773 0·512833 20 
ZIT-99-28 0·703397 12 18·5598 15·5597 38·2240 0·512914 10 
ZIT-99-31 0·703663 10 18·6462 15·5776 38·3761 0·512853 12 
ZIT-99-30 0·704321 12 18·6372 15·5822 38·3509 0·512794 16 
VBZ-HRE        
ZIT-99-25 0·703306 10 18·5713 15·5584 38·2462 0·512978 
ZIT-99-8 0·704355 10 18·6755 15·5879 38·4227 0·512893 18 
ZIT-99-9 0·703847 10 18·6489 15·5787 38·3629 0·512891 12 
ZIT-99–10 0·703351 12 18·6019 15·5627 38·2872 0·512930 14 
ZIT-99-12 0·704253 10 18·6554 15·5782 38·3770 0·512876 14 
ZIT-99-14 0·703692 12 18·6362 15·5710 38·3424 0·512899 18 
ZIT-99-26 0·704217 10 18·6324 15·5759 38·3326 0·512876 12 

2σ errors for individual Sr and Nd measurements are multiplied by 106. Reproducibility for Pb isotopes are given by the double-spike fractionation corrected composition of the NBS981 standard 206Pb/204Pb = 16·9356 ± 0·0048 (143 ppm), 207Pb/204Pb = 15·4912 ± 0·0047 (152 ppm), 208Pb/204Pb = 36·7025 ± 0·014 (191 ppm) (2σ, n = 13). Pb values were adjusted to NBS981 standard values of Todt et al. (1996).

Fig. 2.

Selected major element (a–c) and trace element (d–f) variation diagrams of the studied rock suites. Alkaline and subalkaline division from Irvine & Baragar (1971); K2O vs SiO2 classification fields from Le Maitre (1989). Grey shaded field in (c) indicates experimental melts of natural hydrous basalt at 1–4 GPa (Sen & Dunn, 1994; Rapp, 1995; Rapp & Watson, 1995; Rapp et al., 1999). Calculated molar Mg-number = Mg/(Mg + 0·85Fe2+tot).

Fig. 2.

Selected major element (a–c) and trace element (d–f) variation diagrams of the studied rock suites. Alkaline and subalkaline division from Irvine & Baragar (1971); K2O vs SiO2 classification fields from Le Maitre (1989). Grey shaded field in (c) indicates experimental melts of natural hydrous basalt at 1–4 GPa (Sen & Dunn, 1994; Rapp, 1995; Rapp & Watson, 1995; Rapp et al., 1999). Calculated molar Mg-number = Mg/(Mg + 0·85Fe2+tot).

VBZ trace elements show relative enrichments of the large ion lithophile elements (LILE) over the high field strength elements (HFSE) that are typical of arc lavas (Fig. 3); nonetheless, a close examination of the trace element systematics allows the recognition of two chemically different rock suites, herein named as VBZ high rare earth element (REE) series (VBZ-HRE) and VBZ low REE series (VBZ-LRE). The VBZ-HRE rocks tend to have higher Sr, Ba and REE contents and show distinctively strong Zr–Hf negative anomalies, an unusual feature that has not been previously observed in other TMVB rock suites (Figs 2f and 3b). VBZ-LRE andesites tend to have high Sr/Y ratios at low Y contents (Fig. 4a), a feature often observed in volcanic arcs where young and hot slabs are being subducted, and where a slab-melt component has been interpreted as the subduction agent (Defant & Drummond, 1990; Martin, 1999). VBZ-HRE rocks share the high Sr/Y ratios but they also extend to higher Y contents. Although the VBZ rocks display light and middle rare earth element (LREE–MREE) enrichments over the heavy rare earth elements (HREE) (Fig. 4b), this fractionation is not as strong as observed in ‘adakite-like’ magmas of the Miocene TMVB (Gómez-Tuena et al., 2003) or the Andean Austral Volcanic Zone (Stern & Kilian, 1996); but it is comparable with that in high-Mg andesites from Adak island in the Aleutians (Kay, 1978) and from Mount St. Helens in the Cascades (Smith & Leeman, 1996).

Fig. 3.

N-MORB-normalized trace element patterns of the studied rock suites. (a) Valle de Bravo–Zitácuaro low REE series (VBZ-LRE); (b) Valle de Bravo–Zitácuaro high REE series (VBZ-HRE). VBZ-LRE variation field shown in grey. N-MORB values from Sun & McDonough (1989).

Fig. 3.

N-MORB-normalized trace element patterns of the studied rock suites. (a) Valle de Bravo–Zitácuaro low REE series (VBZ-LRE); (b) Valle de Bravo–Zitácuaro high REE series (VBZ-HRE). VBZ-LRE variation field shown in grey. N-MORB values from Sun & McDonough (1989).

Fig. 4.

Variation of (a) Sr/Y vs Y, (b) Sm/Y vs Sr/Y, (c) Rb/Ba vs Sr/Y, and (d) Cs/La vs K2O/La compared with other Mexican volcanoes such as Popocatépetl (light grey field) at similar fractionation levels. Popocatépetl data from LaGatta (2003) and Schaaf et al. (2005).

Fig. 4.

Variation of (a) Sr/Y vs Y, (b) Sm/Y vs Sr/Y, (c) Rb/Ba vs Sr/Y, and (d) Cs/La vs K2O/La compared with other Mexican volcanoes such as Popocatépetl (light grey field) at similar fractionation levels. Popocatépetl data from LaGatta (2003) and Schaaf et al. (2005).

Relationships between Sr, Nd, and Pb isotope ratios and possible end-member components are shown in Fig. 5. Overall, the Sr–Nd isotopic ratios of the VBZ are bracketed between the compositions of East Pacific Rise mid-ocean ridge basalts (EPR-MORB) and an enriched ‘crustal component’, similar in composition to Mexican lower crustal xenoliths and high-grade metamorphic terranes, or to the bulk subducted sediment sampled at DSDP Site 487. Pb isotopes form a tight positive correlation bracketed by EPR-MORB and our estimate of the average value for Site 487 sediments (which could also represent a crustal component) (Fig. 5c and d). Sr and Pb isotopic compositions largely overlap within the suites, but the Nd isotope compositions of the VBZ-HRE rocks appear to be slightly shifted towards higher values at a given 87Sr/86Sr ratio. Interestingly, rocks from the VBZ tend to overlap the compositions observed at Popocatépetl stratovolcano but they also extend to relatively lower Sr and Pb isotope ratios at similar 143Nd/144Nd ratios.

Fig. 5.

Sr–Nd–Pb isotope variation diagrams for the studied rock suites and possible end-members. (a) 143Nd/144Nd vs 87Sr/ 86Sr; (b, c) Pb isotope variation diagrams; (d) 206Pb/204Pb vs 87Sr/ 86Sr. Also shown: data for 5–15°N East Pacific Rise MORB (MORB) [PETDB, 2006, http://www.petdb.org/(Lehnert et al., 2000)]; weighted bulk sediment composition from DSDP Site 487 (BS) (Gómez-Tuena et al., 2003; LaGatta, 2003); isotopic composition of the altered oceanic crust sampled at DSDP Site 478 (AOC) (Verma, 2000b); Sr–Nd composition of lower crustal xenoliths collected in northern Mexico (LC Xenoliths) (Ruiz et al., 1988a, 1988b; Schaaf et al., 1994); compositional field of ‘intraplate’ basalts (IP-Bas) from the Chichinautzin volcanic field (LaGatta, 2003) and the Palma Sola volcanic field (Gómez-Tuena et al., 2003); compositional field of Popocatépetl stratovolcano (Popo) (LaGatta, 2003; Schaaf et al., 2005).

Fig. 5.

Sr–Nd–Pb isotope variation diagrams for the studied rock suites and possible end-members. (a) 143Nd/144Nd vs 87Sr/ 86Sr; (b, c) Pb isotope variation diagrams; (d) 206Pb/204Pb vs 87Sr/ 86Sr. Also shown: data for 5–15°N East Pacific Rise MORB (MORB) [PETDB, 2006, http://www.petdb.org/(Lehnert et al., 2000)]; weighted bulk sediment composition from DSDP Site 487 (BS) (Gómez-Tuena et al., 2003; LaGatta, 2003); isotopic composition of the altered oceanic crust sampled at DSDP Site 478 (AOC) (Verma, 2000b); Sr–Nd composition of lower crustal xenoliths collected in northern Mexico (LC Xenoliths) (Ruiz et al., 1988a, 1988b; Schaaf et al., 1994); compositional field of ‘intraplate’ basalts (IP-Bas) from the Chichinautzin volcanic field (LaGatta, 2003) and the Palma Sola volcanic field (Gómez-Tuena et al., 2003); compositional field of Popocatépetl stratovolcano (Popo) (LaGatta, 2003; Schaaf et al., 2005).

DISCUSSION

The overall trace element patterns of VBZ lavas are typical of arc-related andesites, for example with Nb–Ta depletions, Pb enrichments, and high Sr/Nd and Ba/LREE (Fig. 3). However, different VBZ rock suites with distinct trace element patterns and isotopic compositions erupted concurrently in the area. In addition, VBZ lavas display compositions that are different from other Mexican stratovolcanoes (Fig. 4). For example, VBZ lavas tend to have higher Sr/Y and Sm/Y ratios, but lower Cs/La and Rb/Ba ratios than Popocatépetl lavas at similar fractionation levels. The geochemical differences observed within the VBZ rocks, and at different geographical positions within the arc, might be related to variations in the subduction geochemical fluxes, or they might be governed by different components and processes that operate during magma ascent through the upper plate (i.e. crystal fractionation, crustal melting and/or contamination). In this section we discuss the different components and processes that controlled the compositions of the VBZ rocks and provide a quantitative model for their petrogenesis.

Crystal fractionation and crustal contamination

Several workers have recognized that a significant part of the geochemical variability observed in continental arcs could be attributed to a process of crustal melting and/or assimilation, such that the chemical imprints derived from the subduction environment are largely obscured (Hildreth & Moorbath, 1988). Indeed, based on isotopic and trace element modeling, some workers have suggested that crustal contamination has played an important role in the petrogenesis of differentiated magmas emplaced across the TMVB (Verma, 1999, 2000a; Lassiter & Luhr, 2001; Chesley et al., 2002; Schaaf et al., 2005). Therefore, evaluating the role of magma differentiation and crustal contamination in the genesis of the studied rocks is important if we aim to recognize and quantify the processes that take place at deeper levels in the subduction zone.

The geochemical trends observed in the major and trace element variation diagrams (Figs 2 and 4), together with the variability in the isotopic compositions of the two VBZ rock suites (Fig. 5), rule out simple crystal fractionation from a common primitive magma as the dominant petrogenetic process. For instance, whereas the major element compositions of both rock suites display coherent trends, Sr and REE contents are significantly higher in the VBZ-HRE than in the VBZ-LRE (Fig. 2). Clearly, these compositional differences cannot be simply accounted for by crystal fractionation from a cogenetic primitive magma, but should instead be a reflection of distinct magma source regions for both rock suites.

Atherton & Petford (1993) showed that some rocks that display slab-melt geochemical features, such as high LREE/HREE and Sr/Y ratios, could be the result of partial fusion of newly underplated, garnet-bearing, basaltic crust. Indeed, several workers have recognized that melting of a thickened crust by underplating of mantle-derived magmas can form SiO2-rich liquids with trace element characteristics that are indistinguishable from those related to partial fusion of the subducted oceanic crust (Conrey et al., 2001; Hou et al., 2004; Wang et al., 2005). The VBZ rocks, however, display significantly higher MgO contents (to ∼9 wt %) and Mg-number (∼75 to 50) that are unlike those attributed to direct partial melting of a mafic garnet-bearing lower continental crust (Mg-number <50, Fig. 2c). In addition, high Sr/Y ratios are coupled with less radiogenic Pb and Sr isotopic compositions (Fig. 6), and thus the isotope ratios of the higher Sr/Y lavas trend towards the composition of Pacific MORB, and away from the isotopic compositions of lower crustal xenoliths and Mexican basement terranes (Fig. 5). Therefore, if the VBZ rocks with high Sr/Y ratios were derived from melting or assimilation of the pre-existent continental crust, the isotopic composition of the contaminant must be more like Pacific MORB than currently known Mexican basement.

Fig. 6.

Relationships between Mg-number, Sr/Y and Pb–Sr isotopic composition for VBZ rocks. Sr/Y ratios correlate inversely with (a) 87Sr/86Sr and (b) 206Pb/204Pb, indicating that the slab-melt geochemical signal tends to have the isotopic characteristics of MORB . (c, d) Variation of 87Sr/86Sr and 206Pb/204Pb vs Mg-number. Calculated molar Mg-number = Mg/(Mg + 0·85Fe2+tot).

Fig. 6.

Relationships between Mg-number, Sr/Y and Pb–Sr isotopic composition for VBZ rocks. Sr/Y ratios correlate inversely with (a) 87Sr/86Sr and (b) 206Pb/204Pb, indicating that the slab-melt geochemical signal tends to have the isotopic characteristics of MORB . (c, d) Variation of 87Sr/86Sr and 206Pb/204Pb vs Mg-number. Calculated molar Mg-number = Mg/(Mg + 0·85Fe2+tot).

Correlations between isotopic compositions and indices of fractionation are often regarded as the best indicators of crustal contamination (assimilation–fractional crystallization; AFC) because this process involves the incorporation of a magma derived from the partial fusion of the crust, and the concurrent fractionation of crystalline phases (DePaolo, 1981). However, the isotopic compositions of the VBZ are largely uncorrelated with Mg-number (Fig. 6c and d), precluding a simple AFC process. Furthermore, VBZ rocks display a puzzling decrease in MREE, HREE (e.g. Fig. 2e), and Nb/Ta ratios (Fig. 7a) with increasing SiO2 contents. Correlations like these are a common feature of many rock suites emplaced throughout the TMVB (Luhr & Carmichael, 1985b; Verma, 1999, 2000b; Straub & Martin-Del Pozzo, 2001; LaGatta, 2003; Siebe et al., 2004; Schaaf et al., 2005). Fractional crystallization, with or without crustal assimilation, that involves the formation of a typical low-pressure mineral assemblage (Ol, Plag, Cpx and Opx) is unable to explain such trends because the HFSE and the REE are not significantly incorporated in these mineral species. Fractionation of Fe–Ti oxides from a basaltic magma will decrease the Nb and Ta contents of the derivative liquids, but it will also increase their Nb/Ta ratios because DNb/DTa < 1 in these phases (Green & Pearson, 1987).

Fig. 7.

Nb/Ta vs SiO2 (a), Dy vs Nb (b) and Nb/Ta vs Nb (c) for VBZ rocks. High-SiO2 andesites from the VBZ tend to have lower Nb/Ta ratios, MREE–HREE and Nb contents than true basalts from the TMVB. These variations are difficult to explain by crystal fractionation (or AFC) of a basaltic magma unless enormous quantities of fractionating amphibole are considered. Fractional crystallization and contamination models (continuous-line and dashed-line arrows, respectively) are shown with different proportions of fractionating amphibole. Models in (b) and (c) use an arc basalt from the Chichinautzin volcanic field as initial magma [sample 297 from Wallace & Carmichael (1999), further analyzed for trace elements by LaGatta (2003), with 10·74 ppm Nb, 0·648 ppm Ta and Nb/Ta = 16·57]. Contamination (AFC) models in (c) use a gneissic enclave (gray square, Xen) from the Chalcatzingo domes (A. Gómez-Tuena, unpublished data) located ∼40 km SW of Popocatépetl volcano as contaminant (Nb 12·13 ppm, Ta 1·22 ppm, Nb/Ta = 9·9). Fractional crystallization models without amphibole use bulk DNb = 0·014, DTa = 0·016 and DDy = 0·167 (assuming fractionation of 35% Ol, 34% Cpx, 20% Opx, 10% Plag, 1% Fe–Ti oxides). Fractional crystallization and AFC (Ma/Mc = 0·8) models involving amphibole use DNb = 0·338, DTa = 0·213 and DDy = 0·277 (assuming 34% Ol, 20% Amph, 20% Opx, 15% Cpx, 10% Plag, 1% Fe–Ti oxides), and DNb = 1· 387, DTa = 0·852 and DDy = 1·235 (assuming 85% Amph, 10% Gt). Also shown in (c) is a simple mixing model (line with black diamonds) of a slab melt (gray diamond, SM) calculated from the composition of the AOC at DSDP Site 487 (Nb 0·49 ppm, Ta 0·034 ppm, Nb/Ta = 14·5; see Table 4 for partition coefficients) and an ‘intraplate’ basalt from the Chichinautzin volcanic field [sample ASC44 from LaGatta (2003) with 24·08 ppm Nb, 1· 427 ppm Ta and Nb/Ta = 16·9]. The analytical precision of the Nb/Ta ratios of repeated standards with similar concentrations to the samples is ∼ 2·5% (2σ), similar to the size of the plotted symbols. Also included in (a) is the EPR-MORB field from Donnelly (2002).

Fig. 7.

Nb/Ta vs SiO2 (a), Dy vs Nb (b) and Nb/Ta vs Nb (c) for VBZ rocks. High-SiO2 andesites from the VBZ tend to have lower Nb/Ta ratios, MREE–HREE and Nb contents than true basalts from the TMVB. These variations are difficult to explain by crystal fractionation (or AFC) of a basaltic magma unless enormous quantities of fractionating amphibole are considered. Fractional crystallization and contamination models (continuous-line and dashed-line arrows, respectively) are shown with different proportions of fractionating amphibole. Models in (b) and (c) use an arc basalt from the Chichinautzin volcanic field as initial magma [sample 297 from Wallace & Carmichael (1999), further analyzed for trace elements by LaGatta (2003), with 10·74 ppm Nb, 0·648 ppm Ta and Nb/Ta = 16·57]. Contamination (AFC) models in (c) use a gneissic enclave (gray square, Xen) from the Chalcatzingo domes (A. Gómez-Tuena, unpublished data) located ∼40 km SW of Popocatépetl volcano as contaminant (Nb 12·13 ppm, Ta 1·22 ppm, Nb/Ta = 9·9). Fractional crystallization models without amphibole use bulk DNb = 0·014, DTa = 0·016 and DDy = 0·167 (assuming fractionation of 35% Ol, 34% Cpx, 20% Opx, 10% Plag, 1% Fe–Ti oxides). Fractional crystallization and AFC (Ma/Mc = 0·8) models involving amphibole use DNb = 0·338, DTa = 0·213 and DDy = 0·277 (assuming 34% Ol, 20% Amph, 20% Opx, 15% Cpx, 10% Plag, 1% Fe–Ti oxides), and DNb = 1· 387, DTa = 0·852 and DDy = 1·235 (assuming 85% Amph, 10% Gt). Also shown in (c) is a simple mixing model (line with black diamonds) of a slab melt (gray diamond, SM) calculated from the composition of the AOC at DSDP Site 487 (Nb 0·49 ppm, Ta 0·034 ppm, Nb/Ta = 14·5; see Table 4 for partition coefficients) and an ‘intraplate’ basalt from the Chichinautzin volcanic field [sample ASC44 from LaGatta (2003) with 24·08 ppm Nb, 1· 427 ppm Ta and Nb/Ta = 16·9]. The analytical precision of the Nb/Ta ratios of repeated standards with similar concentrations to the samples is ∼ 2·5% (2σ), similar to the size of the plotted symbols. Also included in (a) is the EPR-MORB field from Donnelly (2002).

It has also been proposed that amphibole fractionation at shallow levels (Castillo et al., 1999), or fractionation of amphibole and garnet from hydrous basalts at high pressures (Müntener et al., 2001), can explain the slab-melt geochemical features of some adakites (Castillo et al., 1999; Garrison & Davidson, 2003; MacPherson et al., 2006). Amphibole fractionation at either low or high pressures could in principle be responsible for the low MREE, Nb and Nb/Ta ratios in more evolved lavas, but in this case enormous quantities of fractionating amphibole (> 85%) would be required to explain the observed variations (Fig. 7b and c). High-pressure fractionation experiments involving hydrous basalts essentially form residual websterites, with amphibole and garnet appearing at higher degrees of crystallization (Müntener et al., 2001). High-pressure differentiation will also produce corundum-normative derivative liquids (Müntener et al., 2001), and the strongest garnet signatures will probably be observed in the most evolved rocks (MacPherson et al., 2006). Except for one relatively altered sample (ZIT-99-19), none of the VBZ rocks studied here are corundum normative. Furthermore, the strongest garnet signatures are observed in the relatively silica-poor VBZ-HRE (e.g. higher Sm/Y, Fig. 4b), in contrast to what would be expected if both suites were genetically related to deep crystal fractionation. In addition, even if a common parental basalt for the Mexican andesites exists at depth, the trace element and isotopic variations observed within the VBZ suite (Figs 3 and 5), and at different locations along the arc (Fig. 4), will be hard to explain by differentiation alone, and probably require a compositionally heterogeneous primary source.

In conclusion, although some fractionation and interaction with the continental crust is clearly unavoidable, as the magmas must traverse at least 30 km of continental crust before eruption, and there is unequivocal evidence of this interaction in the form of xenocrysts and xenoliths, the overall geochemical variations of the VBZ rock suites cannot easily be accounted for by crystal fractionation from a common primitive magma, nor by crustal assimilation or direct crustal anatexis; therefore, the following discussion assumes that their compositions largely represent mantle-derived and subduction-related phenomena.

Insights into the composition of the Mexican mantle wedge

The concentrations and ratios of the HFSE and the MREE–HREE are often used to infer the characteristics of the mantle source of arc lavas, because these elements are largely insoluble in aqueous fluids and thus are not significantly incorporated into the subduction flux (Class et al., 2000; Hochstaedter et al., 2001; Straub et al., 2004; Pearce et al., 2005). Based on this principle, arc magmas are generally considered to be derived from the partial fusion of a depleted N-MORB-like mantle source. In Mexico the composition of the pre-subduction mantle wedge is elusive because of the compositional heterogeneity of primitive basalts (Luhr & Carmichael, 1985a; Wallace & Carmichael, 1999; Verma, 2000b; Petrone et al., 2003).

Nonetheless, alkaline magmas with negligible subduction signatures and higher Nb contents and Nb/Ta ratios than MORB have been recognized all across the TMVB (Fig. 7). These rocks, often regarded as ‘intraplate’ or ‘ocean island basalt (OIB)-like’ magmas, also have isotopic compositions more ‘enriched’ than MORB (Verma, 2000b; Gómez-Tuena et al., 2003; Petrone et al., 2003). The presence of ‘intraplate-like’ rocks in close space–time association with more typical calc-alkaline arc volcanic rocks has been taken as indicating a highly heterogeneous mantle wedge in the Mexican subduction zone (Luhr & Carmichael, 1985a; Wallace & Carmichael, 1999; Verma, 2000b; Petrone et al., 2003).

Some researchers have suggested that these enriched mantle domains could have been dragged from the back-arc region into the arc source by mantle convection and corner flow (Luhr & Carmichael, 1985a; Wallace & Carmichael, 1999). Others have considered the existence of a mantle plume (Márquez et al., 1999), or association with decompression melting by rifting (Verma, 2002). More recently, Ferrari (2004) suggested that these enriched mantle domains were probably infiltrated into the arc source region following a slab detachment event. However, regardless of the tectonic explanations that have been suggested for the existence of ‘intraplate’ rocks in the Mexican arc, their presence is often considered to be an anomaly that was somehow introduced into the more typical ‘depleted mantle’, because their volume is appreciably small when compared with the more abundant arc-related calc-alkaline sequences.

The decrease of HFSE and MREE–HREE contents, together with Nb/Ta ratios, with increasing SiO2 contents (Fig. 7) is a puzzling characteristic that has been observed in many rock sequences emplaced all across the TMVB (Luhr & Carmichael, 1985b; Verma, 1999, 2000b; Straub & Martin-Del Pozzo, 2001; Siebe et al., 2004; Schaaf et al., 2005). Notably, intraplate volcanic rocks with negligible subduction signals often plot as an enriched end-member in a continuum trend with the rest of the arc-like sequences (Fig. 7). As mentioned above, and as pointed out by previous workers, correlations like these cannot be easily accounted for by a simple liquid line of descent triggered by the crystallization of a typical mineral assemblage, at either low or high pressures, even one extensively modified by crustal assimilation (Fig. 7). In addition, Nb/Ta ratios should not be significantly fractionated by moderate to high extents of mantle melting, because partitioning between these elements and the dominant mantle mineralogy are very similar. Therefore the variations observed in Fig. 7 could, in principle, be attributed to the inherited compositional heterogeneity of the sub-arc mantle wedge beneath Mexico that has been fluxed to various extents by the subduction agents. This interpretation will also mean, however, that rocks with high SiO2 contents may still represent direct mantle melts because Nb/Ta ratios correlate inversely with SiO2, and this variation does not appear to be an artifact of fractionation or crustal contamination.

Water-saturated melting experiments on peridotite have been able to produce high-MgO andesites (Hirose, 1997), and experimental evidence of lavas from the VBZ (Blatter & Carmichael, 2001), together with the presence of amphibole-bearing peridotite nodules in at least one andesitic lava flow (Blatter & Carmichael, 1998a), appears to support the possibility of andesite formation by direct partial melting of the mantle. Nonetheless, to our knowledge, no hydrous melting experiment of peridotite has produced melts with such high silica content as those observed here. The conundrum is thus clear: either there is a mechanism to form very high silica andesites by direct partial melting of the mantle, or the Nb/Ta ratios in this case are simply not reflecting the characteristics of the mantle source.

The usefulness of HFSE ratios as indicators of the mantle source in arc settings relies heavily on the assumption that these elements are not mobilized and/or fractionated by the subduction agents; that is, they behave as ‘conservative’ elements. Support for the insoluble nature of the HFSE in aqueous fluids is clearly established by the experimental evidence (McCulloch & Gamble, 1991; Brenan et al., 1994, 1995c), and the relative low HFSE contents observed in arc rocks with large slab-derived fluid signatures alone (Straub et al., 2004). Nonetheless it has also been increasingly recognized that the HFSE could be mobilized to some extent if the subduction component is in the form of a partial melt coming from the metamorphosed subducted basalt or its sedimentary cover (Stolz et al., 1996; Bourdon et al., 2002). The mobility of these elements, and the possibility to fractionate them, will depend on the residual mineral assembly of the partially melted slab. It has been long considered that rutile should be a residual phase during partial melting of the metamorphosed oceanic crust and sediments, because otherwise the typically high LILE/HFSE ratios observed in modern adakites and Archaean tonalite–trondhjemite–granodiorite (TTG) sequences could not be preserved (Kelemen et al., 2004). Nonetheless, experimental evidence indicates that rutile preferentially incorporates Ta over Nb (DNb/DTa < 1), so that high Nb/Ta ratios might be expected in the partial melts when a rutile-bearing residual assemblage is formed (Green, 1995; Schmidt et al., 2004). In contrast, recently determined partition coefficients between these elements and low-MgO amphiboles indicate that DNb/DTa > 1, and therefore low Nb/Ta ratios might be expected in the slab-derived partial melts when amphibole remains in the residue and the melt fraction is low enough (Tiepolo et al., 2000). Notably, even if some rocks with slab-melt geochemical characteristics have the high Nb/Ta ratios indicative of residual rutile (Stolz et al., 1996; Gómez-Tuena et al., 2003), the vast majority of the studied adakites, as well as the Archean TTG sequences and the bulk continental crust, display subchondritic Nb/Ta ratios (Foley et al., 2002). This characteristic has been taken as evidence for the importance of residual amphibole during partial melting of the subducted slab (Foley et al., 2002), although other investigators have argued that the low Nb/Ta ratios could still be formed with rutile-bearing eclogite residues if the original lithologies involved in the partial melting event had low Nb/Ta ratios to begin with (Rapp et al., 2003).

Although the specific residual mineralogy involved during a slab melting event is hard to define, the experimental results provide insights into the origin and compositional characteristics of the mantle wedge and subduction components involved in the petrogenesis of the Mexican arc. The Nb/Ta variations (Fig. 7) might be explained if the slab-derived chemical flux affecting the arc source region is a water-rich silicate melt derived from the partial fusion of the subducted basalt and sediments, rather than an aqueous fluid in which the HFSE and the HREE are immobile. Low HFSE–HREE contents, together with low Nb/Ta ratios, strong subduction signatures and high SiO2 contents, are all typical characteristics of experimental slab melts and adakites worldwide, and are consistent with melting of garnet amphibolite. These siliceous melts may either react or mix with mantle peridotite and/or their melts during ascent through the mantle and crust (Kelemen, 1998; Prouteau et al., 2001; Tatsumi, 2001). Indeed, Fig. 7c shows that the Nb–Ta concentrations and ratios can be explained by mixing of a slab-derived silicate melt with an enriched ‘intraplate’ magma. The model uses the AOC (altered ocean crust) from DSDP Site 487 as a starting composition (Nb 0·49 ppm, Ta 0·034 ppm, Nb/Ta = 14·5), and assumes garnet amphibolite in the residue (DNb = 0·775 and DTa = 0·555; see Table 4) at 1% batch melting (AOC melt: Nb 0·63 ppm, Ta 0·0607, Nb/Ta = 10·4). Even though very low quantities of residual rutile (< 1%) can significantly affect the Nb/Ta ratios of the resulting slab melt, and sediment contributions are not considered in this case, this exercise provides a feasible explanation for a typical geochemical feature displayed by several volcanic successions of the Mexican arc.

Table 4:

End-member compositions and partition coefficients used in models

 Mantle wedge1 AOC2 Bulk sed.3 DAOC/melt4 Dsed/melt5 Dmantle/melt6 AOC melt7 Sed. melt8 
Rb 0·74 2·94 94·14 0·032 0·72 0·00010 71·11 128·26 
Ba 8·32 8·66 2910 0·025 0·75 0·00011 248·38 3816·79 
Th 0·09 0·10 6·70 0·067 0·89 0·00037 1·35 7·48 
0·027 0·072 3·168 0·066 0·93 0·00110 0·96 3·39 
Nb 1·26 1·77 9·71 0·775 1·21 0·00228 (0·212) 2·27 8·10 
Ta 0·075 0·108 0·673 0·555 1·2 0·00288 (0·243) 0·19 0·57 
La 1·22 2·00 39·31 0·265 1·34 0·00552 7·35 29·71 
Ce 2·81 6·21 48·83 0·380 1·35 0·00819 16·07 36·64 
Sr 30·92 140·00 231·40 0·122 0·51 0·01704 1070·50 432·94 
Pb 0·15 0·26 41·36 0·123 1·29 0·02269 2·01 32·43 
Nd 2·02 6·38 37·68 0·599 1·53 0·02253 10·58 25·06 
Sm 0·54 2·40 7·78 0·924 1·6 0·03672 2·59 4·95 
Zr 14·05 56·85 107·78 0·423 1·64 0·02332 (0·083) 132·74 67·03 
Hf 0·45 1·73 2·42 0·567 1·64 0·03532 (0·095) 3·02 1·50 
Eu 0·20 0·90 2·06 1·400 1·68 0·05000 0·64 1·25 
Gd 0·63 3·65 8·32 1·880 1·65 0·05897 1·95 5·14 
Tb 0·11 0·68 1·31 2·730 1·66 0·07876 0·25 0·81 
Dy 0·65 4·51 7·75 4·140 1·68 0·09666 1·10 4·71 
3·74 27·35 49·15 4·750 1·69 0·11038 5·80 29·69 
Ho 0·13 1·00 1·63 5·355 1·69 0·11667 0·19 0·99 
Er 0·39 2·84 4·55 5·720 1·69 0·13486 0·50 2·75 
Yb 0·37 2·66 4·28 6·770 1·68 0·15196 0·40 2·60 
Lu 0·06 0·41 0·69 9·162 1·72 0·16238 0·04 0·41 
87Sr/86Sr 0·702900 0·702900 0·708458    0·702900 0·708458 
206Pb/204Pb 18·680 18·271 18·687    18·271 18·687 
207Pb/204Pb 15·580 15·489 15·601    15·489 15·601 
208Pb/204Pb 38·400 37·748 38·423    37·748 38·423 
143Nd/144Nd 0·513010 0·513160 0·512518    0·513160 0·512518 
 Mantle wedge1 AOC2 Bulk sed.3 DAOC/melt4 Dsed/melt5 Dmantle/melt6 AOC melt7 Sed. melt8 
Rb 0·74 2·94 94·14 0·032 0·72 0·00010 71·11 128·26 
Ba 8·32 8·66 2910 0·025 0·75 0·00011 248·38 3816·79 
Th 0·09 0·10 6·70 0·067 0·89 0·00037 1·35 7·48 
0·027 0·072 3·168 0·066 0·93 0·00110 0·96 3·39 
Nb 1·26 1·77 9·71 0·775 1·21 0·00228 (0·212) 2·27 8·10 
Ta 0·075 0·108 0·673 0·555 1·2 0·00288 (0·243) 0·19 0·57 
La 1·22 2·00 39·31 0·265 1·34 0·00552 7·35 29·71 
Ce 2·81 6·21 48·83 0·380 1·35 0·00819 16·07 36·64 
Sr 30·92 140·00 231·40 0·122 0·51 0·01704 1070·50 432·94 
Pb 0·15 0·26 41·36 0·123 1·29 0·02269 2·01 32·43 
Nd 2·02 6·38 37·68 0·599 1·53 0·02253 10·58 25·06 
Sm 0·54 2·40 7·78 0·924 1·6 0·03672 2·59 4·95 
Zr 14·05 56·85 107·78 0·423 1·64 0·02332 (0·083) 132·74 67·03 
Hf 0·45 1·73 2·42 0·567 1·64 0·03532 (0·095) 3·02 1·50 
Eu 0·20 0·90 2·06 1·400 1·68 0·05000 0·64 1·25 
Gd 0·63 3·65 8·32 1·880 1·65 0·05897 1·95 5·14 
Tb 0·11 0·68 1·31 2·730 1·66 0·07876 0·25 0·81 
Dy 0·65 4·51 7·75 4·140 1·68 0·09666 1·10 4·71 
3·74 27·35 49·15 4·750 1·69 0·11038 5·80 29·69 
Ho 0·13 1·00 1·63 5·355 1·69 0·11667 0·19 0·99 
Er 0·39 2·84 4·55 5·720 1·69 0·13486 0·50 2·75 
Yb 0·37 2·66 4·28 6·770 1·68 0·15196 0·40 2·60 
Lu 0·06 0·41 0·69 9·162 1·72 0·16238 0·04 0·41 
87Sr/86Sr 0·702900 0·702900 0·708458    0·702900 0·708458 
206Pb/204Pb 18·680 18·271 18·687    18·271 18·687 
207Pb/204Pb 15·580 15·489 15·601    15·489 15·601 
208Pb/204Pb 38·400 37·748 38·423    37·748 38·423 
143Nd/144Nd 0·513010 0·513160 0·512518    0·513160 0·512518 
1

Mantle wedge composition is sample ASC44 from the Chichinatzin volcanic field (LaGatta, 2003) inverted at 5% batch melting to a spinel peridotite (53% Ol, 17% Cpx, 28% Opx, 2% Sp). Individual mineral Kd values from Hart & Dunn (1993), Kelemen et al. (1993), Johnson (1994), Salters & Longhi (1999) and Donnelly (2002).

2

Altered oceanic crust composition is a mixture of 80% AOC from DSDP Site 487 (Verma, 2000b; LaGatta, 2003) and 20% of an average fresh EPR-MORB between 5° and 15°N from the PETDB database (http://www.petdb.org/) with slightly higher Sr contents and 87Sr/86Sr ratios to account for alteration.

3

Bulk sediment composition from DSDP Site 487 from LaGatta (2003) and Gómez-Tuena et al. (2003).

4

Bulk solid–melt partition coefficients of andesitic–dacitic melts in equilibrium with a garnet amphibolite residuum (30% Cpx, 30% Gt, 40% Amph). Individual mineral Kd values from Rollinson (1993, and references therein), Green (1995), Tiepolo et al. (2000), van Westrenen et al. (2001), Barth et al. (2002) and the Geochemical Earth Reference Model (GERM) (http://www.earthref.org).

5

Bulk solid–melt partition coefficients for sediment melting from Johnson & Plank (1999), except for Zr and Hf, which are extrapolations.

6

Bulk solid–melt partition coefficients for a metasomatized mantle wedge assuming a residual mantle mineralogy made of 54% Ol, 10% Cpx, 34% Opx, 2% Gt. Bulk D values for HFSE assuming 0·3% residual rutile are shown in parentheses. Rutile Kd Nb = 70, Kd Ta = 80, Kd Zr = DHf = 20 are best-fit values based on Foley et al. (2000) and Schmidt et al. (2004).

7

Composition of an AOC melt assuming 1% batch melting.

8

Composition of a sediment melt assuming 5% batch melting.

The above-mentioned hypothesis thus implies that the low HREE–HFSE contents and Nb/Ta ratios observed in the silica-rich calc-alkaline rocks could represent the inherited compositional characteristics of a slab-derived silicate melt that had a small contribution from the background mantle wedge (or its partial melts). The greater the interaction of the slab melts with the mantle wedge, or the higher the proportion of the ‘intraplate’ magmatic component in a possible magma mixing event, the larger the dilution effect of the slab-melt component in the resulting magmas. In this scenario, the enriched ‘intraplate’ basalts should best represent the composition of the pre-subduction background mantle wedge in Mexico. This will not imply, of course, that the mantle wedge below Mexico is compositionally homogeneous and identical to an OIB-like source. More research will be needed to test this hypothesis. However, the evidence does indicate that the geochemical tools that have proved useful in other arcs for identifying the composition of the mantle wedge are probably not applicable for the Mexican arc because the HFSE and the MREE–HREE have become an intrinsic ingredient of the subduction flux.

Subduction agents: the effect of AOC and sediment melts

As mentioned above, the VBZ rock suites display contrasting trace element patterns, slightly different isotopic compositions, and often follow diverging trends in several incompatible element ratio plots when compared with other Mexican volcanoes (Figs 3–5). If the geochemical compositions are not significantly controlled by crystal fractionation and/or additions from the continental crust, then one can safely assume that they are controlled by the contributions from the subducted slab and the mantle wedge.

Element ratios between Th, Pb, Nd and Sr are useful indicators of slab-derived chemical fluxes because the partitioning of these elements between slab-derived fluids or melts have been extensively studied (Defant & Drummond, 1990; Drummond & Defant, 1990; Brenan et al., 1995a; Elliott et al., 1997; Johnson & Plank, 1999; Class et al., 2000; Foley et al., 2002; Kessel et al., 2005). Th and Nd are fluid immobile (Brenan et al., 1994, 1995c; Stadler et al., 1998; Kessel et al., 2005), whereas Pb and Sr are soluble in aqueous fluids (Brenan et al., 1994, 1995c; Keppler, 1996; Adam et al., 1997; Ayers et al., 1997; Stadler et al., 1998; Kessel et al., 2005). Therefore, Th and Nd budgets in arc magmas should be mainly controlled by their relative abundances in the mantle wedge plus the slab-melt additions (sediment + AOC). The Pb and Sr contributions, on the other hand, can be controlled by fluids and/or slab-melt additions to the mantle wedge because they are highly soluble, and fairly incompatible during partial melting of eclogite and amphibolite sources (Defant & Drummond, 1990; Tiepolo et al., 2000; van Westrenen et al., 2001; Barth et al., 2002; Foley et al., 2002; Kessel et al., 2005). In addition, because the altered (or fresh) subducted oceanic crust and overlying sediments have different Sr, Nd and Pb isotopic compositions (Verma, 2000b; Gómez-Tuena et al., 2003; LaGatta, 2003), the isotopic variations observed in the magmatic rocks can be used to differentiate between different fluids or melts originating from different portions of the subducted slab.

VBZ samples show a positive correlation between Th/Nd and Pb/Nd, indicating that Pb and Th are incorporated at a constant ratio in the subduction flux (Fig. 8a). The coupling of a fluid-mobile element (Pb) with an insoluble one (Th), and a correlation trending towards higher Th/Nd ratios than those of the subducted sediments, are indications that sediment additions into the mantle source are in the form of a silicate melt rather than an aqueous fluid. This interpretation is further confirmed in the rough but discernible negative correlation of Pb/Nd and Nd isotopes (Fig. 8b), a characteristic that has been often cited as indicating a sediment-melt component that contributes Th, Pb and unradiogenic Nd (Elliott et al., 1997; Class et al., 2000; Hochstaedter et al., 2001; Gómez-Tuena et al., 2003; Kelemen et al., 2004). However, low Pb/Nd ratios extend to lower 143Nd/144Nd values (Fig. 8b) than the EPR-MORB, indicating that the mantle wedge component contributing to the magmas has to be more isotopically enriched (i.e. lower 143Nd/144Nd) than the EPR-MORB source, and more similar to ‘OIB-like’ intraplate basalts.

Fig. 8.

Identification of subduction components in the VBZ rocks. (a) Pb/Nd ratios correlate positively with Th/Nd values, and (b) inversely with 143Nd/144Nd isotopes. These suggest the addition of a sediment partial melt to an isotopically enriched mantle wedge. (c) Sr/Pb ratios define a negative correlation with 206Pb/204Pb that projects towards the composition of the AOC. As Pb is significantly more incompatible than Sr in aqueous fluids, and both elements become similarly incompatible only during partial melting (Kessel et al., 2005), the high Sr/Pb component is consistent with a melt derived from the AOC. (d) The Sr–Pb isotopic variation of the VBZ can be explained by mixing different proportions of melts from the subducted slab (sediments and AOC), but cannot be reproduced if the AOC component is considered to be an aqueous fluid. AOC and bulk sediment compositions, as well as slab-melt partition coefficients, are shown in Table 4. The AOC fluid was calculated assuming 1% fluid release (Sr 48·1 ppm, Pb 0·835 ppm, Sr/Pb = 57·55) using the partition coefficients reported by Kessel et al. (2005) for fluids in equilibrium with an eclogite residuum at 700°C (DSr = 2·93 and DPb = 0·31).

Fig. 8.

Identification of subduction components in the VBZ rocks. (a) Pb/Nd ratios correlate positively with Th/Nd values, and (b) inversely with 143Nd/144Nd isotopes. These suggest the addition of a sediment partial melt to an isotopically enriched mantle wedge. (c) Sr/Pb ratios define a negative correlation with 206Pb/204Pb that projects towards the composition of the AOC. As Pb is significantly more incompatible than Sr in aqueous fluids, and both elements become similarly incompatible only during partial melting (Kessel et al., 2005), the high Sr/Pb component is consistent with a melt derived from the AOC. (d) The Sr–Pb isotopic variation of the VBZ can be explained by mixing different proportions of melts from the subducted slab (sediments and AOC), but cannot be reproduced if the AOC component is considered to be an aqueous fluid. AOC and bulk sediment compositions, as well as slab-melt partition coefficients, are shown in Table 4. The AOC fluid was calculated assuming 1% fluid release (Sr 48·1 ppm, Pb 0·835 ppm, Sr/Pb = 57·55) using the partition coefficients reported by Kessel et al. (2005) for fluids in equilibrium with an eclogite residuum at 700°C (DSr = 2·93 and DPb = 0·31).

The simple scenario of sediment-melt additions into an enriched mantle wedge is further complicated when the contribution of Sr, and the isotopic compositions of Sr and Pb, are taken into account. Figure 8c shows that the VBZ rocks form an inverse correlation between Sr/Pb and Pb isotope ratios that trend toward a DSDP Site 487 sediment end-member with high Pb isotope ratios on one hand, but whose other end-member is an isotopically depleted (low Pb isotope ratio) MORB-like component with Sr/Pb ratios that are higher than EPR-MORB. This component is different from what would be expected from an aqueous fluid derived from the dehydration of AOC, but is consistent with a silicate melt derived from its partial fusion. Although Pb and Sr are soluble elements, experimental data indicate that Pb is significantly more incompatible than Sr in aqueous fluids extracted from a plagioclase-free source, and that it is only when partial melting is achieved that Sr and Pb are incorporated in the melt at similar rate (Brenan et al., 1995b, 1995c; Kessel et al., 2005). Supercritical fluids will preferentially incorporate Sr over Pb, but this will only happen at much higher pressures (Kessel et al., 2005). In addition, Fig. 8d shows that MORB or AOC, bulk Site 487 sediment, and the VBZ data form a hyperbolic trend for Sr–Pb isotope ratios that is consistent with binary mixing. The curvature of the trend can only be reproduced if the depleted MORB-like component is a partial melt with a relatively higher Sr/Pb ratio than that of a fluid extracted from the same source.

Modeling subduction and mantle contributions

Figure 8 shows that the overall compositions of the studied rock suites can be explained by mixing of two chemically distinct subduction components derived from the partial fusion of the subducted sediments and underlying oceanic crust. Participation of an enriched mantle wedge, with lower Nd isotope ratios than EPR-MORB, is also apparent from Fig. 8b. Although the precise chemical compositions of these end-members are difficult to constrain, they can be calculated to a reasonable extent with the geological and compositional information available (Table 4).

We have shown above that the entire trace element budget of the VBZ rocks could have been affected by contributions from the subducted slab. This complicates calculation of the composition of the background mantle wedge in the region. Nonetheless, the data indicate that its composition is more similar to the enriched OIB-like mantle that gives rise to the Mexican intraplate basalts than to the source of EPR-MORB (Figs 7 and 8). For this reason, we calculated the trace element composition of the mantle wedge component by inverting the trace element contents of a MgO-rich basalt from the Chichinautzin volcanic field (Wallace & Carmichael, 1999) that shows very little indication of a subduction-derived component (Table 4). The concentrations of Ba, Pb and Sr in the original sample were slightly reduced to have a smooth trace element pattern in a MORB-normalized trace element variation diagram, but no further corrections were made to account for fractionation. We thus emphasize that the modeled slab contributions to the calculated mantle wedge should be considered as maxima in this case. The Nd isotopic composition of the mantle (143Nd/144Nd = 0·51301) was calculated by extrapolation of the trend observed in Fig. 8b, and is similar to that observed in intraplate basalts from other parts of Mexico (Luhr et al., 1989; Pier et al., 1989; Wallace & Carmichael, 1999; Lassiter & Luhr, 2001; Gómez-Tuena et al., 2003; LaGatta, 2003; Petrone et al., 2003). Because the Pb and Sr isotopic compositions of the studied rock suites are strongly controlled by the slab-derived agents (Fig. 8d), the values for the mantle wedge were assumed to be slightly more enriched than the average EPR-MORB (87Sr/86Sr = 0·7029, 206Pb/204Pb = 18·68, 207Pb/204Pb = 15·58, 208Pb/204Pb = 38·40), and are also similar to the values observed in Mexican intraplate basalts (Table 4).

The composition of the sediment melt was calculated using the weighted bulk sediment composition from DSDP Site 487 (Gómez-Tuena et al., 2003; LaGatta, 2003) and assuming 5% batch melting, using the partition coefficients reported by Johnson & Plank (1999) (Table 4). The models consider no major changes in fluid-mobile element contents during fore-arc devolatilization, and assume that melting of the bulk subducted sediment occurs at arc front depths.

A major uncertainty in this modeling is the composition of the altered oceanic crust. The AOC at DSDP Site 487 (Verma, 2000b) is an obvious possibility; however, these particular samples have extremely depleted REE patterns and higher 143Nd/144Nd ratios when compared with most East Pacific Rise samples [PETDB: Petrological Database of the Ocean Floor, 2006, http://www.petdb.org/(Lehnert et al., 2000)]. Therefore, it is unclear to what extent the Site 487 AOC samples represent the bulk oceanic crust that is being subducted along the Middle American Trench (Gómez-Tuena et al., 2003). Nonetheless, if an average composition of a fresh EPR-MORB is chosen as the starting material, then there will be a large uncertainty on the amount of alteration and mobility of key elements that might be transferred to the arc source region during subduction. For these reasons, the models below use both the composition of the pristine AOC at DSDP Site 487 and a hypothetical AOC calculated by mixing 20% of an average EPR-MORB with the DSDP 487 AOC (Table 4). The compositions of the AOC melts were calculated using the average and/or best-fit partition coefficients of andesitic and dacitic liquids, and assuming 1% batch melting with a garnet amphibolite residuum (Table 4).

Natural and experimental slab melts are not in equilibrium with mantle olivine and should react with mantle peridotite during ascent (Rapp et al., 1999). Various experiments and models indicate that slab-melt–peridotite interaction forms Opx-rich lithologies at the expense of olivine, but other phases have also been identified: Al-augite, garnet, plagioclase and hydrous minerals such as amphibole and/or phlogopite (Kelemen, 1998; Prouteau et al., 1999; Rapp et al., 1999; Tatsumi & Hanyu, 2003; Kelemen et al., 2004). Thus modeling such a reaction strongly depends on the partition coefficients between trace elements and the mineral species that are consumed and created. Some workers have attempted to model these reactions using the AFC formulation of DePaolo (1981), and assuming that the slab melts heat as they rise and decompress while assimilating mantle peridotite (Tatsumi & Hanyu, 2003; Kelemen et al., 2004). Others have modeled the process using a simple mixing approach, considering that small amounts of slab-derived additions will not have a significant effect in modifying the mantle mineral paragenesis (Elliott et al., 1997; Class et al., 2000; Gómez-Tuena et al., 2003). Both kinds of models, although conceptually distinct, tend to provide similar results because the incompatible trace element patterns will be essentially governed by the slab-melt additions and the amount of mantle that could be assimilated or mixed.

Given the large uncertainty in the amount and compositions of the minerals that could be consumed and precipitated during melt–peridotite interaction, we have chosen to use the simple mixing approach. Nonetheless, we acknowledge that the ability for a slab melt to remain liquid during ascent will depend on the so-called ‘effective’ melt/peridotite ratio (Rapp et al., 1999). If this ratio remains high, then slab melts could escape significant modification of their incompatible element contents (Rapp et al., 1999). Relatively high bulk Mg-number and Ni for given SiO2 (see Fig. 2c) will probably be the only reliable indicators of slab-derived andesitic–dacitic melts interacting with the mantle wedge (Kelemen et al., 2003; Kessel et al., 2005). In contrast, if the melt/peridotite ratio is relatively low, then melts could solidify completely and form a modally metasomatized, Opx-rich, hybrid mantle that could be later remelted during decompression and ascent. As the partition coefficients for Opx and Ol and most of the elements considered here are extremely low, then moderate to large extents of melting of such a metasomatized mantle will not significantly change the trace element patterns of the resulting melts, and they will be similar to the composition of the bulk slab component itself, unless the proportions of accessory mineral phases such as amphibole, phlogopite or rutile are significantly large. The absolute concentrations, on the other hand, will be controlled by the slab-melt/peridotite ratio, the residual mineralogy, and the extent of melting.

Modeling results

Figure 9 shows modeling results for slab-melt and mantle interactions in Sr, Nd and Pb isotopic variation diagrams. If a melt of the AOC sampled at DSDP Site 487 is used as the depleted MORB-like component, the mixing line in Sr–Nd isotopic space (Fig. 9a) passes over the isotopic compositions of Popocatépetl volcano but is largely unable to reproduce the VBZ data. This probably indicates that the AOC component contributing to VBZ has a slightly different composition from the one contributing to Popocatépetl. For this reason, we slightly adapted the AOC isotopic compositions to fit the observed VBZ geochemical trends, while keeping the bulk-sediment constituent fixed. The isotopic values for this new ‘AOC’ are within the range of those observed in EPR-MORB for Pb and Nd but have a slightly higher 87Sr/86Sr ratio and Sr contents to account for alteration. The rest of the trace elements represent a mixture of 80% AOC from DSDP Site 487 and 20% of an average fresh sample from the EPR (Table 4).

Fig. 9.

Simple mixing modeling results of slab-melt–mantle interactions for the studied rock suites in (a) Sr–Nd and (b) Pb–Nd isotopic space. A mixing model that uses the DSDP 487 AOC can explain, to a large extent, the compositions of the volcanic products of Popocatépetl stratovolcano (Popo) but is unable to explain the compositions of the VBZ suite. Thus a slightly different AOC was constructed to account for the discrepancy (Table 4). The VBZ-LRE rocks are interpreted to be slab melts with little mantle contribution, and with isotopic compositions that reflect variable proportions of sediment and AOC melts. An average VBZ-LRE (large filled triangle) can be formed by a mixture of ∼ 70% AOC melt and ∼30% sediment melt. (c) and (d) show modeling results of slab-melt additions to the mantle wedge (MW). Adding ∼10% of a bulk slab component (made of 60% AOC and 40% sediment melt) into the calculated mantle wedge can closely reproduce the isotopic composition of an average VBZ-HRE rock (▵).

Fig. 9.

Simple mixing modeling results of slab-melt–mantle interactions for the studied rock suites in (a) Sr–Nd and (b) Pb–Nd isotopic space. A mixing model that uses the DSDP 487 AOC can explain, to a large extent, the compositions of the volcanic products of Popocatépetl stratovolcano (Popo) but is unable to explain the compositions of the VBZ suite. Thus a slightly different AOC was constructed to account for the discrepancy (Table 4). The VBZ-LRE rocks are interpreted to be slab melts with little mantle contribution, and with isotopic compositions that reflect variable proportions of sediment and AOC melts. An average VBZ-LRE (large filled triangle) can be formed by a mixture of ∼ 70% AOC melt and ∼30% sediment melt. (c) and (d) show modeling results of slab-melt additions to the mantle wedge (MW). Adding ∼10% of a bulk slab component (made of 60% AOC and 40% sediment melt) into the calculated mantle wedge can closely reproduce the isotopic composition of an average VBZ-HRE rock (▵).

By using this hypothetical AOC the VBZ-LRE rock suite samples plot within a mixture made of 80:20 and 60:40 AOC melt:sediment melt (Fig. 9). An average VBZ-LRE rock could be explained by a mixture of about 70:30 AOC melt:sediment melt (large filled triangle in Fig. 9), with very little contribution from the mantle wedge. The trace element pattern of such a mixture is very similar to an average VBZ-LRE rock, except for the fluid-mobile elements (Rb, Ba, U and Pb) (Fig. 10a). Several possible explanations can account for these differences. They might reflect the above-mentioned uncertainty in the composition of the AOC that is being subducted, or it could be that some of these elements have been retained in hidden accessory mineral phases that were not considered in the models, such as residual mica [e.g. Elburg et al. (2002)]. Nonetheless, the remarkable match in the rest of the elements, and the fact that the discrepancy affects only the most fluid-mobile elements, makes these interpretations unlikely. We consider it a more plausible explanation that a portion of these highly soluble elements has been simply lost in a fluid phase prior to melting at arc-front depths. This is conceivable because the solidus temperatures of wet basalt or sediments are significantly higher than the stability field of several of the hydrated mineral species that become unstable at relatively low metamorphic grades during subduction (Johnson & Plank, 1999; Poli & Schmidt, 2002). Thus, it seems inevitable that fluid release from the subducted slab should precede anatexis. As the slab component contributing to VBZ-LRE is clearly a composite, it is difficult to quantify precisely the extent of element retention or depletion that affected the sediments and/or the altered oceanic crust. Nevertheless, if the modeled VBZ lavas are true liquid compositions, as previously indicated by the experimental evidence (Blatter & Carmichael, 2001), or have undergone a small amount of fractionation or trace element dilution during mantle interaction, then ∼ 66% of the Rb, ∼ 65% of the Ba, ∼ 50% of the U and ∼40% of the Pb must have been removed (or retained) from the available fluid-mobile element budget before being incorporated into the calculated bulk-slab component.

Fig. 10.

(a) Model N-MORB-normalized trace element pattern compared with an average VBZ-LRE rock. The trace element pattern of an average VBZ-LRE rock is closely matched by a mixture of 70% AOC melt and 30% sediment melt, except for the fluid mobile elements. This suggests that these elements were either retained in an accessory residual phase or, more probably, released in a fluid phase before melting was achieved. (b) Model N-MORB-normalized trace element patterns show that most of the trace elements of an average VBZ-HRE can be reproduced by batch melting (∼7% F) of a mantle wedge that has been metasomatized with ∼10% slab melt composed of 60% AOC and 40% sediment, except for the fluid-mobile elements and the HFSE. Even though the fluid-mobile element difference can be explained by retention or fluid release, the strong negative Nb–Ta and Zr–Hf anomalies probably necessitate the participation of ∼0·3% residual rutile in the mantle wedge (see text and Table 4 for modeling parameters).

Fig. 10.

(a) Model N-MORB-normalized trace element pattern compared with an average VBZ-LRE rock. The trace element pattern of an average VBZ-LRE rock is closely matched by a mixture of 70% AOC melt and 30% sediment melt, except for the fluid mobile elements. This suggests that these elements were either retained in an accessory residual phase or, more probably, released in a fluid phase before melting was achieved. (b) Model N-MORB-normalized trace element patterns show that most of the trace elements of an average VBZ-HRE can be reproduced by batch melting (∼7% F) of a mantle wedge that has been metasomatized with ∼10% slab melt composed of 60% AOC and 40% sediment, except for the fluid-mobile elements and the HFSE. Even though the fluid-mobile element difference can be explained by retention or fluid release, the strong negative Nb–Ta and Zr–Hf anomalies probably necessitate the participation of ∼0·3% residual rutile in the mantle wedge (see text and Table 4 for modeling parameters).

The VBZ-HRE rocks tend to have higher MgO and Ni contents, more radiogenic Nd isotope compositions, and lower Ba/La, Th/Nd and Pb/Nd ratios than the VBZ-LRE suite, features that are all indicative of a relatively higher contribution from the mantle wedge (Fig. 8). Figures 9c and d shows the effects of mixing a bulk-slab component (made of 60:40 AOC melt:sediment melt) into the calculated mantle wedge. The proportion of slab component needed to reproduce the isotopic compositions of an average VBZ-HRE rock can be constrained to be ∼10%. Batch melting of such a metasomatized mantle (90% mantle: 10% slab-melt component) can reproduce fairly well the overall trace element patterns of an average VBZ-HRE rock if the fluid-mobile element contents of the bulk slab are adapted for depletion (or retention) in the proportions described above (Fig. 10b). Nevertheless, a model that assumes a peridotitic (or even pyroxenitic) residual mantle mineralogy is unable to reproduce the strong negative Nb–Ta and Zr–Hf anomalies of the VBZ-HRE rock suite because the partitioning of the HFSE and their neighboring elements in the normalized trace element diagrams are considered to be very similar. The HFSE should remain unaffected during the fluid loss that could have affected the LILE because these elements are insoluble in aqueous fluids. In addition, even though negative Zr–Hf anomalies are a common feature of sediments at DSDP Site 487, and are especially strong in the lower pelagic (non-terrigenous) section of the column, these sediments also have MORB-like Pb isotopic compositions (LaGatta, 2003), and their preferential incorporation into the VBZ-HRE magmas should be evident in plots involving Pb isotopes. As the VBZ-LRE and VBZ-HRE rocks are practically indistinguishable in their Pb isotopic compositions, the origin of the Zr–Hf anomalies observed in the VBZ-HRE cannot be related to the preferential incorporation of DSDP Site 487 pelagic sediments.

The presence of residual amphibole and/or rutile is often considered to play a significant role in the formation of Nb–Ta anomalies in slab-derived melts, but very little has been written about the effect on Zr and Hf. As mentioned before, the low Nb/Ta contents and ratios of the VBZ-LRE probably derive from the effect of residual amphibole during slab melting, but this same mineral has also been used to explain the absence of fractionation between Zr–Hf and the REE, because low-Mg amphibole favors the MREE over Zr and Hf (Foley et al., 2002). At first sight it might appear that a different residual mineralogy involving rutile (or zircon) during slab melting explains the negative Zr–Hf anomalies observed in the VBZ-HRE suite. In detail, however, this is not sufficient to elucidate why these rocks appear to have a higher contribution from the mantle wedge, or why those more likely representatives of pristine slab melts (the VBZ-LRE) do not display negative Zr–Hf anomalies.

A possible solution to this dilemma is that the Zr–Hf anomalies have been formed during partial fusion of a parcel of mantle that has been intensively modified by the addition of slab-derived silicate melts. Indeed, if slab melting does take place, then strong reaction of these melts with the surrounding mantle will occur unless the melt/peridotite ratio remains high (Rapp et al., 1999). If the effective melt/peridotite ratio of an average VBZ-HRE magma is assumed to be as low as that constrained by the isotope diagrams (90% mantle: 10% slab melt), then it is very likely that the slab-derived melts were readily consumed and transformed into pyroxene-rich metasomatic lithologies. If these veins were later remelted in the presence of a residual accessory phase (e.g. rutile or zircon), the resulting magmas will display trace element patterns with unusual anomalies in those elements that became more compatible. In fact, Fig. 10b shows that a very small amount of residual rutile (∼0·3%) during partial melting of such a modally metasomatized mantle can account for the strong Nb–Ta and Zr–Hf negative anomalies that are characteristic of the VBZ-HRE rocks. This could also explain the relatively high Nb/Ta ratios of the VBZ-HRE suite at relatively low Nb contents (Fig. 7).

A question that remains is whether an accessory titaniferous phase such as rutile can precipitate and remain stable in the mantle wedge after a partial melting event. Experimental studies and calculations have shown that rutile is not a likely residual phase in the source of basaltic melts, as the TiO2 contents in basalts are too low (Green & Pearson, 1986; Ryerson & Watson, 1987); however, other studies have indicated that rutile could become stable in peridotite–melt systems under hydrous, silica-rich conditions (Schiano et al., 1994, 1995; Bodinier et al., 1995). In addition, accessory rutiles have also been directly observed and analyzed in modally metasomatized peridotites from Siberia (Ionov et al., 1999; Kalfoun et al., 2002). These studies have shown that rutile could become an important HFSE-bearing phase at least in the conditions of the upper mantle, and thus provide independent support for the presence of residual rutile during partial melting of a modally metasomatized mantle wedge.

CONCLUSIONS AND IMPLICATIONS

The geochemical data and modeling results provided here indicate that slab melting plays a key role in the petrogenesis of andesitic rocks of Valle de Bravo–Zitácuaro, in the central Trans-Mexican Volcanic Belt. Although these magmas have traversed a 30–40 km thick continental crust, the geochemical results do not support simple crystal fractionation or crustal contamination as the main petrogenetic processes. Instead, the geochemical variations within the studied rock suites can be largely attributed to contribution of partial melts derived from the subducted basalt and its overlying veneer of sediment interacting with the mantle wedge. The new geochemical evidence thus supports previous experimental studies that suggested a mantle origin for andesites in the TMVB (Blatter & Carmichael, 2001; Carmichael, 2002). Nonetheless, the new data further suggest that a slab-derived siliceous melt interacting with the mantle wedge is probably a more effective mechanism to form high-SiO2 andesites than water-saturated partial melting of the upper mantle.

The recognition that slab melting plays an important role in the Mexican arc, together with the conspicuous presence all across Mexico of ‘intraplate’ magmas with enriched (low Nd, high Sr, Pb) isotope ratios, yields a new perspective on the chemical character of the Mexican mantle wedge. Indeed, the results presented here indicate that the HFSE do not remain immobile during slab melting, and thus their concentrations and ratios in subduction-related lavas do not necessarily resemble those of the mantle source. Instead, the depleted mantle signatures (i.e. the low Nb/Ta ratios), together with strong subduction signals and high SiO2 contents, could reflect the inherited character of the slab-melt component itself.

The models provided here also support results from experimental petrology (Johnson & Plank, 1999; Poli & Schmidt, 2002) indicating that the composition of subducted slabs changes during their journey into the mantle, and that they undergo intense modifications of their water content and fluid-mobile element budget as a result of continuous fluid release. In turn, we can distinguish the effects of different mechanisms of element recycling (hydrous fluids vs silicate melts) in the geochemistry of arc magmas because they imprint different geochemical characteristics on the volcanic products. Furthermore, because these differences should be ultimately linked to variations in PT gradients, they provide an insight into the thermal and tectonic structure of a subduction zone. For instance, the low Nb/Ta ratios of the VBZ-LRE rocks indicate that slab melting probably occurred in the presence of amphibole, independently constraining slab depth to ∼75 km, equivalent to < 25 kbar, because amphibole is unstable at higher pressures (Schmidt & Poli, 1998). If true, then it is possible that the Cocos slab remains flat for a longer distance, and melts at a significantly shallower depth than the hypothetical ∼100 km below the central part of the arc (Pardo & Suárez, 1995). A detailed seismic tomography of central Mexico would be capable of testing this hypothesis.

Slab-melt–peridotite reactions also appear to be viable mechanisms for the formation of accessory mineral assemblages in the mantle wedge. As silicate melts are about 10 times more effective as metasomatic agents to form hydrous minerals than aqueous fluids (Schneider & Eggler, 1986), it does not seem surprising that the only reported locality in Mexico with amphibole-bearing lherzolites is the VBZ (Blatter & Carmichael, 1998a), and that their host is a VBZ-HRE rock (our sample ZIT-99-10). The presence of residual accessory phases such as amphibole, phlogopite or even rutile not only can selectively imprint unusual geochemical signals on the arc magmas, but could also represent important carriers of key trace elements in the convecting mantle to the temperatures and pressures at which they remain stable.

ACKNOWLEDGEMENTS

A.G.-T. thanks Alex LaGatta for fruitful discussions on the petrogenesis of the TMVB. Invaluable help during sample preparation and analysis was provided by Rick Mortlock, Marty Fleisher, Jean Hanley and Rufino Lozano. We thank Marlina Elburg and Colin MacPherson for reviews that led to significant improvements in the manuscript. Editorial handling by Professor John Gamble is also highly appreciated. This work was funded by CONACyT grants 39785 to A.G.-T. and 27642-T to F.O.-G.; and by the NSF grant EAR 96-14782 to C.H.L. and S.L.G.

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