Southward migration of volcanic activity in the central Mexican Volcanic Belt: asymmetric extension within a two-layer crustal stretching model

Journal of Volcanology and Geothermal Research 112 (2001) 175±187 www.elsevier.com/locate/jvolgeores

Southward migration of volcanic activity in the central Mexican Volcanic Belt: asymmetric extension within a two-layer crustal stretching model Alvaro MaÂrquez a,*, Roberto Oyarzun b, Cristina de Ignacio a, Miguel Doblas c a b

Departamento PetrologõÂa y GeoquõÂmica, Facultad de Ciencias GeoloÂgicas, Universidad Complutense, 28040 Madrid, Spain Departamento CristalografõÂa y MineralogõÂa, Facultad Ciencias GeoloÂgicas, Universidad Complutense, 28040 Madrid, Spain c Museo Nacional de Ciencias Naturales (CSIC), Jose Gutierrez Abascal 2, 28006 Madrid, Spain Received 27 November 2000; revised 15 May 2001; accepted 15 May 2001

Abstract Magmatic activity in the Miocene to present Mexican Volcanic Belt (MVB) is migrating to the south. Volcanic activity is associated with extension, which is being accommodated along a major, 1000-km-long, E±W-oriented continental rift. In turn, the MVB is cross-cut by older, though active N±S- to NNW-oriented major extensional faults, along which the polygenetic volcanoes are aligned. A southward migration pattern is also observed for the monogenetic volcanism, which however appears to be related to E±W- and N608E-oriented extensional faults. We suggest that the southward migration of volcanic activity can be explained in terms of a two-layer crustal stretching model (brittle and ductile domains). The layers would be separated at an upper crustal level by a zone of simple-shear decoupling, at the brittle±ductile transition zone (BDTZ). The overall movement above the BDTZ is southward-directed, the only direction to which the Central MVB can extend and grow. Our model suggests that the magmas that feed the volcanism become stored at the BDTZ. Evidence supporting this assumption is provided by the Al-in-hornblende thermobarometer and the Ol±Pl±SiOr pseudoternary diagram, which indicate average pressures between 2.5 and 4 kbar. The magmas feeding the monogenetic volcanism ascend rapidly along active E±W and N608E extensional faults (large strain rates), i.e. they do not have enough time to form large magma chambers. The magmas feeding the polygenetic volcanism are emplaced along N±S to NNW faults (lower strain rates). These magmas remain stored for longer periods at the BDTZ of the N±S to NNW faults, and therefore form large magmatic chambers, shaping vertical overshoots of several kilometers of height. The results from geobarometry indicating magma emplacement depth at around 8 km for the polygenetic volcanism, and 12 km for the monogenetic volcanism, are in good agreement with the rheological constraints of a BDTZ at about 10 km of depth. We envisage a feedback mechanism regarding magma storage and shallowing of the BDTZ, i.e. magma emplacement shallows the BDTZ; in turn, this shallowing controls the new zones for magma emplacement, a southward directed process. q 2001 Elsevier Science B.V. All rights reserved. Keywords: Mexican Volcanic Belt; migration of volcanism; upper crust; extension; decoupling

1. Introduction * Corresponding author. Present address: ESCET, Universidad Rey Juan Carlos, C/Tulipan s/n 28933 Mostoles Madrid, Spain. Tel.: 134-91-4887017; fax: 134-91-6647490. E-mail address: [email protected] (A. MaÂrquez).

The 1000-km-long, E±W-oriented Mexican Volcanic Belt (MVB) is an active volcanic province crossing southern Mexico. The MVB

0377-0273/01/$ - see front matter q 2001 Elsevier Science B.V. All rights reserved. PII: S 0377-027 3(01)00240-2

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Fig. 1. The MVB, depicting the active volcanic centers (triangles) and main monogenetic volcanic ®elds (Chichinautzin and MichoacaÂnGuanajuato) (based on Nixon et al., 1987).

displays a clear pattern of southward directed migration of volcanic activity from Mio±Pliocene time onwards (Fig. 1). Mio±Pliocene volcanism is located at the northern realm of the belt, whereas active volcanism including the great stratovolcanoes (e.g. PopocateÂpetl) and the recent monogenetic scoria cones (e.g. ParicutõÂn and Jorullo, MichoacaÂn±Guanajuato volcanic ®eld; e.g. Luhr and Carmichael, 1985), is located in the southern front of the belt. We have concentrated on the central part of the Mexican Volcanic Belt (CMVB, Figs. 1 and 2) because this is a well known area, where polygenetic and monogenetic volcanism is migrating to the south. The tectonic setting is complex and involves the interplay of NNW and E±W active fault systems, both of them with associated seismic and volcanic activity. The study area is delimited to the west by the Queretaro±Taxco Fault System (Fig. 2), a NNW- trending fault system related to an outstanding gravimetric anomaly. The area is bound to the east by the Sierra de RõÂo FrõÂo±IztaccõÂhuatl±PopocateÂpetl alignment (Fig. 2). To the east of this volcanic alignment the relationships between tectonic activity and volcanism are less evident. Age of polygenetic volcanic activity decreases to the south along well-de®ned NNW volcanic trends.

Although less conspicuous, the southward migration of magmatism is also observed in the monogenetic volcanic ®elds (Cuenca de MeÂxico and Sierra Chichinautzin; Nixon et al., 1987) (Fig. 2). No clear explanation has been yet offered for such migration. As noted by Nixon et al. (1987), the gradual migration of volcanism towards the volcanic front in the Quaternary cannot be related to documented plate reorganizations, and occurred much too rapidly to represent a direct response to a signi®cant change in dip of the subducted slab. The same would apply to trench retreating phenomena (e.g. Singh and Pardo, 1993). Furthermore, to induce a rapid, subduction-related southward migration of magmatism in the MVB, the Cocos plate should be subjected to major rollback type phenomena (e.g. Gvirtzman and Nur, 1999), which does not seem to be the case. Additionally, the geochemical evidence suggest that at least part of the recent volcanism of the southern front of the CMVB is unrelated to the subduction of the Cocos plate (Verma, 1999, 2000a). In this work we concentrate on the southward migration of volcanism in the CMVB, and propose a model based on a two-layer crustal stretching process.

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Fig. 2. (a) Shaded relief map (digital elevation model) for the central Mexican volcanic belt (CMVB) (see inset in Fig. 1); UTM coordinates, zone 14. Large arrows indicate main migration vectors for polygenetic volcanism; Ma: age of volcanoes in million years (after Cantagrel et al., 1981; Nixon et al., 1987; Mora-Alvarez et al., 1991; Ferrari et al., 1999; Osete et al., 2000); 1±3: trends as in text. (b) Main structural features of the Palo HueÂrfano-Nevado de Toluca alignment (1), see the E±W and NNW systems of normal faults; EP: early Pliocene, LM: late Miocene; LP: late Pliocene, Q: Quaternary (simpli®ed after Alaniz-Alvarez et al., 1998).

2. The Mexican Volcanic Belt: a peculiar magmatic arc 2.1. General Many of the geological, geophysical, and petrological characteristics of the MVB make it a peculiar case of a magmatic arc (e.g. MaÂrquez et al., 1999a). The belt is oriented at a 15±208 angle to the Middle America trench, so that the trench is some 400 km from the eastern end of the MVB, but only about 150 km from the western end. In our study area (CMVB; Figs. 1 and 2) the Wadati±Benioff zone is absent below the volcanic front (Pardo and SuaÂrez,

1995), which is curious for a not-so-young subducting plate (20±24 Ma under the MVB; Nixon, 1982). For example, seismicity of Nazca plate under the magmatic arc is recognized in southern Chile (central and southernmost southern volcanic zones: CSVZ± SSVZ; as de®ned by LoÂpez-Escobar et al., 1995), where the age of subducting Nazca (Cifuentes, 1989) is in the same order as Cocos under the MVB. Additionally, there is a remarkable gravity low (, 2 200 mGal Bouguer anomaly) beneath the CMVB (Molina-Garza and Urrutia-Fucugauchi, 1993). Fix (1975) suggested the existence of a lowdensity mantle layer and the presence of partially molten material at the base of the lower crust. The

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local peculiarity of the CMVB nor a time-restricted phenomenon. They occur throughout the MVB and range in age from Miocene to Quaternary. For a discussion on these topics in the CMVB see for example: Ferrari and Rosas-Elguera (1999), MaÂrquez et al. (1999a,b,c), Wallace and Carmichael (1999) and Verma (2000a,b). 2.2. Extensional processes

Fig. 3. Seismic activity in the CMVB (1990±1998; M . 3). Distribution of tectonic epicenters in the CMVB and southern Mexico (from Servicio SismoloÂgico Nacional de MeÂxico). I: IztaccõÂhuatl, NT: Nevado de Toluca, P: PopocateÂpetl, SC: Sierra de Las Cruces. UTM coordinates, zone 14.

latter is interpreted to have Vp ˆ 7:6 km s21 ; and density ˆ 3:29 g cm23 (Molina-Garza and UrrutiaFucugauchi, 1993). The model of Fix (1975) is further supported by data from Gomberg and Masters (1988), indicating very low shear velocities in the lower lithosphere beneath the MVB region. Additionally, they ®nd an S-wave shadow zone, indicative of the presence of a mantle shear wave low-velocity zone that extends to about 300 km depth. Last but not least, geochemically distinct types (dominant andesites±dacites and minor OIB-type ma®c rocks) of volcanic rocks occur in close association (MaÂrquez et al., 1999a,b; Verma, 1999, 2000a). Although much could be argued either in favor or against a subduction origin for the calc-alkaline suite, the OIB magmatism still remains to be explained within a subduction scenario. These basalts are neither a

The geology of central and southern Mexico is characterized by the presence of a series of suspect terranes displaying contrasting lithologic and structural characteristics (Campa and Coney, 1983; Sedlock et al., 1993), and the E±W-oriented MVB, stretching for some 1000 km across central Mexico (Fig. 1). Although most of the attention has been traditionally focused onto the MVB activity, it should be noted that a very important volcanism also developed in the southern terranes from Paleocene to Miocene time (e.g. MoraÂn-Zenteno et al., 1999). Although from Middle Miocene onwards the volcanic activity had ceased to the south of the MVB, the extensional conditions have prevailed (Meschede et al., 1997). Gravimetric data (Campos-EnrõÂquez et al., 2000) show a sharp vertical passage from the older northern units (Sierra Madre Oriental) to the CMVB. This passage is indicative of an E±W-oriented structural discontinuity bounding the CMVB to the north. In the same way, the southern boundary of the CMVB is characterized by a dramatic decrease in altitude from 2200 m at the plateau to 1300 m in the Cuernavaca valley (Fig. 2). In that area gravimetric data (Bouguer anomaly) vary from 2200 (southern sector of the CMVB) to 2150 mGal in the south (Sierra Madre del Sur) in less than 100 km (Molina-Garza and Urrutia-Fucugauchi, 1993). Additionally, the southern realm of the CMVB displays important seismic activity associated with NNW and E±W tectovolcanic structures (Fig. 3). Focal mechanisms indicate E±W-oriented normal faults with a signi®cant strike-slip sinistral component (UNAM-CENAPRED Seismology Group, 1995). Volcanic activity has concentrated along the MVB from Miocene time to the present. Data supporting a generalized extensional scenario throughout the MVB are provided by pervasive E±W normal faults (see for example Fig. 2) displaying seismic activity (Suter et

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al., 1995a). To this respect, MaÂrquez et al. (1999a) propose that the MVB can be regarded as a propagating rift, opening from west to east. This interpretation is supported by structural and tectonic data from the CMVB. Suter et al. (1995a) indicate a rotational deformation component in the CMVB, with a pole of rotation being located to the east of the zone of deformation. The propagating character of the MVB rifting process is supported by the geometry of regional graben-type structures, with a wide western sector and a narrow eastern end (scissors-like; e.g. Acambay graben). Furthermore, the northern boundaries of these graben record NNE-directed extension (with a minor dextral component), whereas the southern borders show SSE-directed extension and a major sinistral component (e.g. Suter et al., 1995a,b; GarcõÂa-Palomo et al., 2000). This implies that extension rates increase to the west. Last but not least, a series of presently active N±S- to NNW-striking normal faults (Fig. 2), related to the southern continuation of the US Basin and Range (Henry and Aranda-Gomez, 1992) complete this complex tectonic scenario. 2.3. Migration of volcanic activity in the CMVB As noted by many authors (Cantagrel and Robin, 1979; Nixon et al., 1987; Alaniz-Alvarez et al., 1998; Osete et al., 2000), the polygenetic and monogenetic volcanism in the CMVB is migrating to the south. In the northern sector of the CMVB the volcanic rocks of the Late Miocene stratovolcanoes (Palo Huerfano, La Joya) (Fig. 2) are located above the Ma®c Volcanic Unit (as de®ned by Ferrari et al., 1994; Henry and Aranda-Gomez, 2000). This ma®c Miocene volcanic sequence (which extends beneath younger rocks throughout the belt, and marks the onset of volcanism in the MVB), is a large and massive volcanic episode that took place between 11 and 6 Ma (Ferrari et al., 1994). At the central area of the CMVB (around Acambay graben) the volcanic activity concentrates in the Pliocene (e.g. Amealco caldera activity; 4.7± 2.2 Ma; Aguirre DõÂaz, 1996) (Fig. 2), with minor Pleistocene activity (Suter et al., 1995a). Finally, at the southern front of the CMVB, the Quaternary volcanic rocks of Nevado de Toluca (a polygenetic volcano) and Sierra Chichinautzin volcanic ®eld (monogenetic volcanoes) (Fig. 2) are directly located

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above the Ma®c Volcanic Unit (GarcõÂa-Palomo et al., 2000), with no evidence of Late Miocene or Pliocene MVB volcanism. Therefore, it can be concluded that by Late Miocene (i.e. after the emplacement of the Ma®c Volcanic Unit), the CMVB volcanism was concentrated in its northern sector, reaching the central area by Pliocene time, and the southern front in the Quaternary. The most striking evidence for the migration of volcanism in the CMVB is provided by a series of NNW- to N±S-oriented polygenetic volcanic alignments (names listed from north to south) (Fig. 2): (1) San Miguel de Allende±Taxco (187 km); Palo HueÂrfano volcano (12.1 Ma), La Joya volcano Ä ado volcano (10.6 Ma), Amealco caldera (4.7 Ma), N (3.43 Ma), Nevado de Toluca volcano (1.5 Ma to present); (2) Sierra de Las Cruces (80 km); Nativitas (3.71 Ma), Monte Alto (2.90 Ma), Villa Alpina (2.87 Ma), Desierto de los Leones (1.92 Ma); and (3) Sierra Nevada (19 km); IztaccõÂhuatl (0.9± 0.27 Ma), PopocateÂpetl (0.27MIN to present). There, the ages of volcanic activity (Cantagrel et al., 1981; Nixon et al., 1987; Mora Alvarez et al., 1991; AlanizAlvarez et al., 1998; Ferrari et al., 1999; Osete et al., 2000) decrease southward, clearly de®ning migration vectors. Migration velocities along these trends deduced from age data indicate ®gures in the following order: (1) Palo Huerfano±Nevado de Toluca: 21 km Ma 21; (2) Sierra de las Cruces: 21 km Ma 21; (3) IztaccõÂhuatl±PopocateÂpetl: 27 km Ma 21. Regarding monogenetic activity, evidence for southward migration is provided by paleomagnetic data for the Quaternary Mexico Valley and Sierra Chichinautzin volcanic ®elds (Mooser et al., 1974; Fig. 2). Whereas the Mexico Valley monogenetic volcanoes display reverse polarity, i.e. Matuyama epoch (.0.78 Ma), those of the Sierra Chichinautzin ®eld (some 30 km to the south) show normal polarity (Brunhes epoch). Furthermore, 14C data for the latter indicate ages younger than 40,000 years (Bloom®eld, 1975).

3. Discussion 3.1. General Alaniz-Alvarez et al. (1998) indicate that the

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polygenetic volcanoes are associated with N±S to NW faults, whereas the monogenetic volcanoes are associated with E±W faults. While the E±W extensional trend in the volcanic areas is well de®ned (e.g. Suter et al., 1995a; MaÂrquez et al., 1999b; GarcõÂaPalomo et al., 2000), the N±S to NNW tectonic trends are less conspicuous to some authors (e.g. Siebe et al., 1999; Suter, 1999), who indicate that major faults with recent activity are not observed in the vicinity of the Quaternary polygenetic volcanoes. However, as noted by Alaniz-Alvarez et al. (1999), this would be the consequence of magmatism accommodating extension along low-strain zones. Thus, the structural trends occupied by magma bodies (dikes, magma chambers) can accommodate strain without faulting. In any case, fault activity is well documented along important N±S to NNW fault systems in the CMVB such as at the San Miguel de Allende±Taxco system (south of QuereÂtaro segment; Alaniz-Alvarez et al., 1998,1999; at Nevado de Toluca area; GarcõÂa-Palomo et al., 2000) or at the Perales fault (Fig. 2). Note that the major volcanic centers (e.g. La Joya Amealco) (Fig. 2) are strictly related to NNW-striking extensional faults. In turn, these faults can be ascribed to the southern continuation of the US Basin and Range in northern and central Mexico (e.g. Henry and Aranda-Gomez, 1992). Although the alignment of polygenetic volcanoes at speci®c trends (e.g. Sierra Nevada, Sierra de Las Cruces) is now well documented and beyond dispute (see for example Fig. 2), the actual mechanism for the southward migration still remains to be explained. To estimate the amount of extension in the CMVB, we followed the area-balance method described by Groshong (1994). We used a well-de®ned regional reference level (top of the Mesozoic limestone unit underlying the volcanic rocks) from a 200-km-long N±S gravimetrical cross-section in the CMVB (Campos-EnrõÂquez et al., 2000) and a depth of detachment of 10 km. Based on these procedures, an estimation of 8.5% of extension (minimum) can be made (i.e. some 17 km) which agrees with values calculated for continental rift systems (less than 20%; Baldridge et al., 1995). A simple-shear extensional scenario, involving the elevation of the Moho at the intersection between the detachment surfaces and the upper mantle, and large-scale crustal thinning (e.g. Wernicke, 1985) does not ®t the CMVB case.

Geophysical data (Urrutia-Fucugauchi and FloresRuiz, 1996) suggest that the Moho topography under the CMVB is irregular (no clear pattern can be deduced), varying between 50 and 45 km under the studied alignments of volcanoes. These data have to be taken, nevertheless, with much caution, as the gravimetric model (Molina-Garza and Urrutia Fucugauchi, 1993) implies a lower crust with a high density of 3.29 g cm 23. This value is closer to those of layers of anomalous mantle beneath areas undergoing continental rifting than to those of `typical' lower crust (in the order 2.8 g cm 23; e.g. Braile et al., 1995). Based on the latter, and on the crustal models for gravity pro®les under the CMVB (Molina-Garza and Urrutia Fucugauchi, 1993), a Moho depth estimate can be made in the order of 40 km. Alternatively, the values from Molina-Garza and Urrutia-Fucugauchi (1993) could be indicative of the existence of a `solid lower crust' (as de®ned by Wernicke (1990)), with Vp in the order of 6.8± 7.4 km s 21, and a density of 2.9±3.1 g cm 23. However, rheological and petrological constraints to be discussed later, make this assumption doubtful. As discussed above, no large-scale crustal extension is observed in the CMVB. For example, contrary to the US Basin and Range, no core-complex-type structural features are found within the MVB. In the same way, other processes suggested for the Basin and Range such as lower crustal material ¯owed laterally to even out the topography of the Moho, from areas of less extension to the areas of greater extension (e.g. Block and Royden, 1990; Wernicke, 1990) can hardly be applied to the Mexican case. If lower crust were laterally ¯owing towards zones of greater extension, then southern Mexico should not have such a remarkable thin crust (35±20 km) (Urrutia-Fucugauchi, 1986). Alternatively, if extension in southern Mexico is regarded as a two-stage scenario, involving the unroo®ng of core-complexes (e.g. the Xolapa complex, along the Paci®c coast; Robinson et al. 1989) and magmatic activity (volcanism between the Xolapa complex and the MVB; MoraÂn-Zenteno et al., 1999) from Eocene to Miocene time, then, we might regard this zone as a highly extended realm formed by a presently cooled ¯uid layer (e.g. Wernicke, 1990). Since the crust in the south is thin (25±30 km; Urrutia-Fucugauchi, 1986), we must admit then that the process either aborted before

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Fig. 4. Rheology (a), density (b), and integrated effective buoyancy (c), for the crust beneath the CMVB. BDTZ is at ,10 km where magmas become stored.

reaching its peak, or the southern realm is presently behaving as a rigid block (`cooled ¯uid layer'), undergoing further extension and therefore crustal thinning. This assumption is supported by structural evidence indicating that the major northward dipping mylonite surface (detachment horizon) bounding the Xolapa complex to the north, is cut by active normal faults, which also tilt to the north blocks of Miocene volcanic rocks. Furthermore, these brittle faults also cut ¯uvial deposits, supportnig a presently active extensional ®eld affecting southern Mexico (e.g. Ratsbacher et al., 1991; Meschede et al., 1997). In any case, a scenario of such a kind would imply two contrasted realms. A southern zone undergoing major extension and magmatism during Eocene±Miocene, and a northern front (the future MVB) with much smaller rates of extension, and magmatic activity starting in Middle Miocene time (Ma®c Volcanic Unit; Ferrari et al., 1994). Whatever be the case, we must regard the continental crust as mechanically and rheologically strati®ed, with crustal levels weaker and less rigid than others, which ultimately allows the development of decoupling phenomena (Axen et al., 1998). Geobarometric and rheological constraints to be discussed later in this section, suggest that the model that better ®ts the present CMVB is a two-layer crustal model. An open system, two-layer crustal stretching model (Gans, 1987; Gans et al., 1989) may account for

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most the structural, petrological, and geophysical features of the CMVB. The lack of large-scale crustal thinning in the CMVB receives a satisfactory explanation in terms of the addition of new materials to the lower crust (open system), which would compensate high-level crustal extension, i.e. while the upper crust is being thinned, the addition of magmas (from a mantle source) to the lower crust keep crustal thickness constant. Evidence supporting this scenario is provided by the existence of the already discussed low-density mantle layer (Vp ˆ 7:6 km s21 ; 23 density ˆ 3:29 g cm ; Molina-Garza and UrrutiaFucugauchi, 1993). We regard this layer as the probable source of ma®c magmatism (e.g. basalts from Sierra Chichinautzin; Verma, 2000a) part of it being added to the lower crust. Note, that a rheological change boundary (the Moho in this case) acts as preferential magma emplacement site (Watanabe et al., 1999). Additionally, this mantle layer can be interpreted as a heat engine for at least part of the felsic magmatism in the CMVB. This is supported by geochemical and isotopic evidence suggesting that the evolved magmas from Sierra Chichinautzin (at the front of the belt) originated from partial melting of a heterogeneous ma®c granulite source in the lower crust (Verma, 1999). A key element for a two-layer stretching model is the existence of a BDTZ acting as a decoupling surface between a heterogeneously deforming (brittle) upper crust from a uniformly deforming (ductile) mid and lower crust (Gans, 1987). In order to characterize this boundary beneath the CMVB we constructed a rheological pro®le, which indicates a BDTZ at 10 km of depth (Fig. 4a). We adopted a steady state geotherm with a surface heat ¯ow of 90 mW m 22 (Ziagos et al., 1985), a strain rate of 10 215 s 21, a quartz rheology for the upper crust (0±20 km), plagioclase rheology for the lower crust (20±40 km), and wet dunite rheology for the upper mantle (Ranalli, 1997). 3.2. Constraints for magma emplacement The emplacement of magmas is expected at the level of neutral buoyancy (LNB), i.e. the level at which the density difference between magma and host rocks changes its sign (Fig. 4b). The differential stress can signi®cantly affect the ascent and

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Fig. 5. Plotting of bulk compositions of samples from polygenetic and monogenetic hornblende dacites from the CMVB, from normative magnetite 1 diopside onto olivine±plagioclase± (quartz 1 orthoclase) diagram (Ol±Pl±SiOr) (Baker, 1987). The 1 atm and 8 kbar liquid lines of descent (dashed lines) are dry and the 2 and 5 kbar ones (solid lines) are for compositions with 2 wt% H2O. Squares are samples from IztaccõÂhuatl volcano (Nixon, 1988b) and open circles from Sierra Chichinautzin monogenetic volcanoes (Verma, 1999; Wallace and Carmichael, 1999).

emplacement of magmas. Therefore the LNB usually occurs at rheological boundaries. To make an estimate of the depth of magma emplacement beneath the CMVB, we have calculated the integrated effective buoyancy (Fig. 4c) by subtracting the differential stress (Fig. 4a) to the LNB (e.g. Watanabe et al., 1999). In the CMVB the integrated effective buoyancy is maximized at the Moho and at the BDTZ (Fig. 4c) where magmas will be preferentially emplaced. Evidence supporting the emplacement of magmas at the BDTZ (10 km) in the CMVB is provided by estimations of crystallization pressures of hornblende dacites from the IztaccõÂhuatl polygenetic volcano (Nixon, 1988a), and from the Sierra Chichinautzin monogenetic volcanoes (Verma, 1999; Wallace and Carmichael, 1999). These samples contain the assemblage Bt 1 Qtz 1 Pl 1 Mt ^ Ilm; which is suitable for the application of several geobarometers. Because of the lack of analytical data for amphiboles from the Sierra Chichinautzin rocks, we used the Ol±Pl±SiOr pseudoternary diagram of Baker (1987). This geobarometer was calibrated for high-alumina basalts, andesites, and dacites. The experimentally produced melts, saturated with Pl 1 anhydrous ma®c minerals

(Ol, augite, pigeonite and orthopyroxene) were projected on pseudoternary diagrams to de®ne different lines of multiple saturation as a function of pressure and water content in the magmas. The experiments were run at dry conditions (curves of 1 atm and 8 kbar) and at hydrous (2% H2O) conditions (curves of 2 and 5 kbar) (Baker and Eggler, 1987). Hornblende dacites from the CMVB plot close to the 2 kbar curve (Fig. 5). This particular system (Ol±SiOr±Pl), is appropiated because of the ubiquitous occurrence of diopside and magnetite in the lavas under consideration. The obtained crystallization pressure was tested with the conventional Al-in-hornblende thermobarometer (Johnson and Rutherford, 1989). The most recent calibration of this thermobarometer after Anderson and Smith (1995) cannot be applied to the samples as these authors discouraged the use of their barometer for hornblendes with Fetot =…Fetot 1 Mg† outside their calibration range (0.4±0.65). The Fetot =…Fetot 2 Mg† ratios of the hornblendes from IztaccõÂhuatl and Sierra Chichinautzin samples vary from 0.22±0.36 (Nixon, 1988b; MaÂrquez, 1999). The Schmidt (1992) calibration does not ®t our requirements because of the following reasons: (1) the temperature range is much lower than ours (655±7008C); and (2) the hornblendes used for the calibration incorporate 80% of their Al content in tschermakitic-type substitutions, whereas in the IztaccõÂhuatl samples, the main substitution is of the edenite-type. On the contrary, the geobarometer of Johnson and Rutherford (1989) is calibrated for a higher temperature range (up to 7808C) which is roughly in agreement with the minimum temperatures (8208C) obtained for IztaccõÂhuatl samples from Mt± Ilm pairs (Nixon, 1988b). In addition, the Johnson and Rutherford (1989) geobarometer was applied to volcanic rhyodacites and rhyolites with mineral assemblages similar to those from IztaccõÂhuatl. Although the mineral assemblage required by the geobarometer is not complete in our samples (lacking sanidine and sphene), Johnson and Rutherford (1989) state that the lack of sphene is likely to have a negligible effect on pressure estimates. Furthermore, Anderson and Schmidt (1995) indicate that the absence of sphene and potassium feldspar does not appear to signi®cantly affect pressure determinations. The calculated pressures for the IztaccõÂhuatl and Sierra Chichinautzin hornblende-dacites are shown

A. MaÂrquez et al. / Journal of Volcanology and Geothermal Research 112 (2001) 175±187 Table 1 Pressure values obtained for hornblende-dacites from Iztac: IztaccõÂhuatl polygenetic volcano CHI: Sierra Chichinautzin) with the Alin-hornblende barometer of Johnson and Rutherford (1989) Sample

Total Al content

P (^0.5 kbar)

Iztac.: Core R3 Iztac.: Core R3 Iztac.: Core R7 Iztac.: Rim R3 Iztac.: Rim R3 Iztac.: Rim R7 CHI-49: Core

1.4370 1.3997 1.3845 1.2118 1.2161 1.3845 1.7600

2.62 2.46 2.40 1.67 1.68 1.27 3.98

in Table 1. The hornblende cores of the IztaccõÂhuatl samples yield crystallization pressures around 2:5 ^ 0:5 kbar; well in agreement with those obtained by the Baker (1987) pseudoternary diagram. The hornblende rims of these samples, however, show slightly lower pressure values, due to the fact that they have partially undergone syn-eruptive oxidation. The only hornblende analysis available from Sierra Chichinautzin dacites yields a rather higher pressure value of 3:98 ^ 0:5 kbar (around 12 km). Despite the monogenetic character of the volcanism of Sierra Chichinautzin, which would imply the absence of large crustal magmatic reservoirs, the mineralogical and textural features of the studied dacites (e.g.

Fig. 6. Earthquake distribution in depth. The arrow indicate the depth (,15 km) above which 80% of the earthquakes have occurred. Same data set as in Fig. 3.

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phenocryst-rich, large and complex zoned plagioclase; MaÂrquez, 1999; MaÂrquez et al., 1999b) are in agreement with an emplacement of magmas in the upper crust (at the BDTZ) beneath the volcanic ®eld. To this respect, the structural and geodynamic setting of the Sierra Chichinatuzin monogenetic volcanism (see MaÂrquez et al., 1999b) explains both the monogenetic character of the volcanism and the greater emplacement depth of its magmas. Since these magmas ascend rapidly along the active E±W and N608E extensional faults (large strain rates; AlanizAlvarez et al., 1998), they do not have enough time to form large magmatic chambers. In other words, the magmas undergo a short-lived storage. On the contrary, the magmas feeding the polygenetic volcanism are emplaced along NNW faults (lower strain rates). These remain stored for longer periods at the BDTZ of the NNW faults, and therefore form large magmatic chambers, shaping vertical overshoots of several kilometers of height (e.g. Watanabe et al., 1999). Therefore, the results from geobarometry indicating magma emplacement depth at around 8 km for the polygenetic volcanism and 12 km for the monogenetic volcanism, are in good agreement with the rheological constraints of a BDTZ at about 10 km of depth. Thermal modelling of calderas (temperature ®eld distribution) in the study area (Los Azufres, MichoacaÂn) (Fig. 2) as well as outside it (Los Humeros Puebla; La Primavera Jalisco) suggest depths of magma chambers in the order of 5±10 km (Verma et al., 1990; Verma and Andaverde, 1996; Verma and RodrõÂguez-GonzaÂlez, 1997). Seismicity constraints for a 10 km deep BDTZ are dif®cult to evaluate, because the Mexican seismic network is not dense enough, which implies that the error in depth estimates can be in the order of several kilometers (J. Pacheco, Head of the Servicio SismoloÂgico Nacional de MeÂxico; personal communication, 2000). The model of crustal seismic wavespeed used to determinate earthquake location has an interface at 16 km of depth. Seimic activity beneath the MVB continues below our 10 km estimate for the BDTZ and a sharp decrease of seismic activity occurs at around 15 km (data obtained upon request; Servicio SismoloÂgico Nacional de MeÂxico) (Fig. 6). This discrepancy (10 versus 15 km depth) can be a function of the model of crustal seismic wavespeed used (interface at 16 km).

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Fig. 7. a and b: Highly schematic cross-sections representing for simplicity a given N±S extensional trend. From a to b extension and shallowing of the BDTZ; decoupling zone causes the progressive southward diversion of magmas up to their ®nal storage beneath the BDTZ. From there upwards magmas ascend along brittle normal faults giving rise to volcanic activity. LDUM: low-density upper mantle.

4. Proposed model Our model for the southward directed migration of volcanic activity in the CMVB involves the interplay of the following processes (Fig. 7): (1) a continuous input of ma®c magmas into the lowermost crust from the underlying anomalous mantle layer; those ascending along N±S to NNW extensional faults (low-strain faults) will accumulate at the BDTZ, whereas those ascending along E±W extensional faults (high-strain faults) will more rapidly reach the surface; in other words, the latter will remain for shorter periods at the BDTZ; (2) generalized extension from the CMVB to the southern coastal Ranges (Sierra Madre del Sur) (Fig. 1); and (3)a BDTZ shallowing to the south in response to magma accumulation. Therefore, we propose that one of the key elements for the understanding of the southward migration of the CMVB volcanism is the tectonic setting of the southern realm (MVB to the coast). Southern Mexico has been subjected to extension for the last 20 Ma (Meschede et al., 1997), and the main directions re¯ect both the E±W and N±S to NNW structural trends. The BDTZ may play an important role in the migration of volcanism. We suggest that this horizon

represents more than a simple rheological change, because it is precisely along this transition where the ascending magmas are stored prior to their ®nal ascent along brittle normal faults (e.g. Gans, 1987; Gans et al., 1989; Watanabe et al., 1999). Additionally, as indicated by Kusznir and Park (1987) and Park (1988), zones of low ductile strength will form decoupling horizons located at high/mid crustal levels in extensional orogens subjected to high geothermal gradients (80±90 mW m 22). Thus, the BDTZ can be regarded as both, a magma storage zone, and an active decoupling horizon, allowing southward-directed extension of the CMVB. If the BDTZ shallows to the south then the magmas should show a tendency to migrate towards that direction. This is because magmas will move and accumulate in those zones of progressively lower P regimes. We obviously envisage a feedback mechanism regarding magma storage and shallowing of the BDTZ, i.e. magma emplacement shallows the BDTZ (Fig. 7); in turn, this shallowing controls the new zones for magma emplacement (Fig. 7). On the other hand, the southward-directed extensional process is a direct consequence of the southern extensional ®eld, i.e. these lowlands, characterized by a

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low-stress ®eld are the only direction to which the CMVB can extend and grow. This process may be further enhanced because of the following: a topographic load causes collapse and expansion of elevated mountains towards low-altitude forelands (the average vertical stress is larger beneath the mountains) (e.g. Burg and Ford, 1997). Finally, upper crustal southward-directed extension must be horizontally accommodated along a major decoupling horizon, which in the Mexican case may be constituted by the BDTZ, at an estimated depth of about 10 km (Figs. 4 and 7). 5. Conclusions The southward directed migration of volcanism in the CMVB is a process controlled by magma accumulation beneath a BDTZ. As noted by Watanabe et al. (1999), in extensional areas the tectonic stress dominates the integrated effective buoyancy, therefore the rheological boundaries primarily control the emplacement of magmas through the effect of tectonic stresses. Additionally, the CMVB brittle±ductile zone behaves as an upper crustal decoupling horizon, with top-to-the-south movement (Fig. 7). We are obviously aware that the actual geometry of the interacting systems may be far more complicated than what is shown in the schematic model depicted in Fig. 7, and may include widespread anastomosing fault/ shear zones along which brittle and ductile conditions are met. As shown in the same ®gure our model predicts the southward shallowing of the BDTZ (decoupling zone), which would result in the migration of volcanism in the same direction, i.e. magmas will tend to accumulate in those progressively shallower zones of the BDTZ (Fig. 7b), which in turn will further shallow the BDTZ, in other words, a feedback mechanism. Acknowledgements This work was partly funded by a grant from Universidad Complutense (Madrid Spain)-Universidad Nacional AutoÂnoma de MeÂxico (Bilateral Program) Projects CONACYT 0196P-T DGAPA IN-106199 (MeÂxico) and DGES PB-95-0107 (Spain). We thank Francisco Anguita for long and

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helpful discussions on the Mexican Volcanic Belt. We also thank Javier Pacheco, Head of the Servicio SismoloÂgico Nacional de MeÂxico, for providing the authors with the data set on Mexican earthquakes that we used in this paper, and Teresa Mateos and Bartolo Luque (Centro de AstrobiologõÂa; INTA-CSIC) for helping with the rheological pro®les. We thank Joann Stock and Surendra P. Verma for constructive comments on the paper. References Aguirre-DõÂaz, J., 1996. Volcanic stratigraphy of the Amealco caldera and vicinity, Central Mexican Volcanic Belt. Rev. Mex. Cienc. Geol. 13, 10±51. Alaniz-Alvarez, S., Nieto-Samaniego, A.F., Ferrari, L., 1998. Effect of strain rate in the distribution of monogenetic and polygenetic volcanism in the Transmexican volcanic belt. Geology 26, 591±594. Alaniz-Alvarez, S., Nieto-Samaniego, A.F., Ferrari, L., 1999. Effect of strain rate in the distribution of monogenetic and polygenetic volcanism in the Transmexican volcanic belt: reply. Geology 27, 573±575. Anderson, J.L., Smith, D.R., 1995. The effects of temperature and fO2 on the Al-in-hornblende barometer. Am. Mineral. 80, 549±559. Axen, G.J., Selverstone, J., Byrne, T., Fletcher, J.M., 1998. If the strong crust leads, will the weak crust follow? GSA Today 8 (12), 1±8. Baker, D.R., 1987. Depths and water contents of magma chambers in the Aleutian and Mariana island arcs. Geology 15, 496±499. Baker, D.R., Eggler, D.H., 1987.Compositions of anhydrous and hydrous melts coexisting with plagioclase, augite, and olivine or low-Ca pyroxene from 1 atm to 8 kbar: application to the Aleutian volcanic center of Atka, 72, pp. 12±28. Baldridge, W.S., Keller, G.R., Braile, L.W., 1995. Continental rifting: a ®nal perspective. In: Olsen, K.H. (Ed.), Continental Rifts: Evolution, Structure, Tectonics. Developments in Geotectonics 25. Elsevier, Amsterdam, pp. 453±459. Block, L., Royden, L.H., 1990. Core-complex geometries and regional scale ¯ow in the lower crust. Tectonics 9, 557±567. Bloom®eld, K., 1975. A late-Quaternary monogenetic volcano ®eld in central Mexico. Geol. Rundschau 64, 476±497. Braile, L.W., Keller, G.R., Wendlandt, R.F., Morgan, P., Khan, M.A., 1995. The East African rift system. In: Olsen, K.H. (Ed.). Continental Rifts: Evolution, Structure, Tectonics. Developments in Geotectonics 25. Elsevier, Amsterdam, pp. 325± 344. Burg, J.P., Ford, M., 1997. Orogeny through time: an overview. Orogeny Through Time, Burg, J.P., Ford, M. (Eds.). Geol. Soc. Spec. Publ. 121, 1±7. Campa, M.F., Coney, P.J., 1983. Tectono-stratigraphy terranes and mineral resource distributions in Mexico. Can. J. Earth. Sci. 20, 1040±1051.

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