no flat wadati–benioff zone in the central and southern central andes

www.elsevier.com/locate/tecto

Tectonophysics 395

No flat Wadati–Benioff Zone in the central and southern

central AndesB

Miguel Munoz*

International Centre for Theoretical Physics, Strada Costiera 11, 34014 Trieste, Italy

Received 27 November 2002; accepted 6 September 2004

Available online 2 December 2004

Abstract

The Wadati–Benioff Zone (WBZ) is an approximate plane defined by earthquakes hypocentres observed in convergent plate

boundaries and that usually dips at angles greater than 308. In some areas of the Andes, where there are gaps in volcanic activity,

and where heat flow is abnormally low, this plane in most studies has nearly horizontal dip at a depth of about 75–100 km, and

it has been associated to flat subduction of the oceanic lithosphere. This situation has been taken as the present-day analogue of

the Laramide orogeny of western North America for which a dflat-slabT episode has been proposed in the past years. In this

work, the observed low heat flow in areas of the Andes is assumed to be due to low radiogenic heat generation in geologically

old and allochthonous terranes constituting large regions of western South America. On the basis of geotherms obtained for

areas of Ecuador, Peru, Chile and Argentina, and of rheological results describing the partition between brittle and ductile

regimes, the seismic activity observed both in the lower crust and at depths of about 75–100 km is thoroughly explained. At

these depths, earthquakes occur within the subcontinental upper mantle, and then there is no flat WBZ associated to subduction

of the oceanic lithosphere. There is evidence from recent seismological observations that the real WBZ lies not horizontally and

deeper in the tectonosphere.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Flat Wadati–Benioff Zones; Andes; Heat flow; Rheology; Earthquakes

1. Introduction

The spatial distribution of earthquakes defining the

Wadati–Benioff Zone (WBZ) beneath western South

America usually results in cross sections that show

0040-1951/$ - see front matter D 2004 Elsevier B.V. All rights reserved.

doi:10.1016/j.tecto.2004.09.002

B Paper written as visiting scientist at ICTP Trieste.

* Present address: Jorge Matte 2005, Santiago, Chile.

E-mail address: [email protected].

segments of inclined seismic zones, with dip angles of

about 308 beneath several areas (Barazangi and

Isacks, 1979; Cahill and Isacks, 1992; Engdahl et

al., 1998; Taboada et al., 2000). In areas of Ecuador

(18N–28S), in northern and central Peru (58S–158S)and in central Chile and western Argentina (27.58S–338S), the WBZ as described by several authors has

nearly horizontal dip (e.g., Barazangi and Isacks,

1979; Cahill and Isacks, 1992; Engdahl et al., 1998).

(2005) 41–65

M. Munoz / Tectonophysics 395 (2005) 41–6542

The flat segments of the WBZ are in correlation with

gaps in Quaternary volcanism observed in Peru, Chile

and Argentina (Fig. 1), although considerable volcan-

ism of Miocene and Pliocene time has been observed

in these areas (Barazangi and Isacks, 1979; Cahill and

Isacks, 1992; Kay and Abbruzzi, 1996; Kay and

Mpodozis, 2002). In Ecuador (Fig. 2), the modern

adakitic arc magmatism has been related to flat-slab

subduction (Gutscher et al., 2000a,b). In Argentina,

between 308S and 328S, the cessation of arc and back-

arc volcanism has been associated to the achievement

of flatness of the subducting Nazca plate in the last 20

Ma (Kay and Abbruzzi, 1996; Kay and Mpodozis,

2002). This would be the late stage of flat-slab

subduction, a process that in Ecuador should yet not

Fig. 1. Contours to depth of the Wadati–Benioff Zone (Cahill and

Isacks, 1992). Epicentres of teleseismic occurrences used to

determine the contours are shown by small crosses. A separation

of the depth contours in Peru between about 58S to 158S, and in

Chile and Argentina from 27.58S to 338S, describes what has beennamed as flat segments of the WBZ.

be fully accomplished (Gutscher et al., 2000a). Such a

distinct process has been assumed to be related to

changing patterns of volcanism, geochemical signa-

tures, subduction of the along-strike prolongations of

the Nazca and Juan Fernandez Ridges, and to

tectonophysical processes as uplift, crustal shortening

and thickening, and formation of foreland basins in

western Argentina (Jordan and Allmendinger, 1986;

Kay and Abbruzzi, 1996; Gutscher et al., 2000b; Kay

and Mpodozis, 2002). However, the assumption of a

flat WBZ (e.g., Kay and Abbruzzi, 1996; Gutscher et

al., 2000a) presents several paradoxes and difficult

problems in connection with the former tectonophys-

ical processes. The suggestion that crust needs to be

lost at about 308S—possibly by fore-arc tectonic

erosion in the continental margin to the west—

depends critically on poor constrained average mod-

ern and Early Miocene crustal thicknesses (Kay and

Abbruzzi, 1996). Also, it has been argued that cooling

of the asthenospheric wedge as the slab becomes flat

causes the removal of the deepest part of the crust

under the Principal Cordillera along with lithospheric

mantle to accommodate the flattening of the sub-

duction zone by creating gravitational unstable con-

ditions (Kay and Kay, 1993; Kay and Abbruzzi,

1996). An objection to this argument is that standard

determination of the WBZ (see Cahill and Isacks,

1992; Smalley et al., 1993) shows that beneath the

Principal Cordillera it dips at angles similar to those

determined north and south of the dflat-slabT segment,

and that dflatteningT is observed from beneath the

Argentine Precordillera terrane at about 708W to the

east (Fig. 1). A dynamical model of thinning of the

lithosphere proposed by Yuen and Fleitout (1985) has

been invoked in general terms by Kay and Kay (1993)

and Kay and Abbruzzi (1996) for explaining the

mechanical removal of the lithosphere, but it should

be noted that such model corresponds to a process

associated to a hot spot, where at the end of the

process the material at the lithosphere–asthenosphere

boundary becoming ductile enough to creep plunges

into the underlying asthenosphere, thereby eroding the

lithosphere. Moreover, for resolving problems related

to geochemical signatures observed in magmas and to

basement uplift of the central Chile–Argentina dflat-slabT segment, it has been assumed that the depth of

the brittle–ductile transition in the crust changed

rather extensively beneath the zones of active volcan-

Fig. 2. Shallow subduction inferred from teleseismic occurrences and the supposed thermal structure (isotherms in 8C) beneath areas of Ecuadorexplaining the active badakiticQ volcanic arc (Gutscher et al., 2000a). This would represent an early stage in the process of slab shallowing whichin central Chile and Argentina is seen to be more evolved as no Quaternary volcanic activity is there observed. Instead, the real Wadati–Benioff

Zones (WBZ) in Ecuador as determined by Guillier et al. (2001) through intensive local seismological studies are planes dipping at angles

between about 308 and 458. Consequently, the badakiticQ arc magmatism is not due to melting of a flat slab.

M. Munoz / Tectonophysics 395 (2005) 41–65 43

ism during the magmatic evolution of the area, thus

allowing the translation of brittle upper and lower

crust into ductile lower crust beneath the volcanic

zones in a intracrustal mixing process associated with

Fig. 3. Simplified map of Ecuador with main morphological units. Sutures o

et al., 2001). The shallow seismicity (0–25 km) is concentrated beneath the

fault (B). There is a lack of shallow seismicity between the trench and the w

line area. Beneath the Andes, seismicity at intermediate depths (25–75 km)

the locations of suture 3 and fault B, with parallel linear patterns that cut

crustal thickening (Kay and Abbruzzi, 1996), or

implying thermal weakening of the crust leading to

thick-skinned basement uplift (Ramos et al., 2002).

This is in conflict with the present-day observed effect

f accreted terranes and main Andean faults are shown (after Guillier

Andean Cordillera, especially between the suture (3) and the reverse

estern slope of the Andes, except for a strip restricted to the coastal

is bounded by two east-dipping lines, which intersect the surface at

the surface at the locations of sutures 2 and 1.

Fig. 4. Geological map of Peru (after Bellido, 1969; Myers, 1975)

The tectonic framework during Cretaceous and early Tertiary time

related to vertical block faulting is also shown. 1, Oceanic crust; 2

Paracas Geanticline; 3 and 4, West Peruvian Trough; 5, Maranon

Geanticline; 6, East Peruvian Trough; 7, Brazilian Shield. In

northern and central Peru, most of the deep seismicity in the crus

is observed east of the western border of the Precambrian basemen

area for about 200–300 km. The dflatT surface of seimicity at 100–

120 km depth lies mainly below the zones of enhanced crusta

seismicity. Active volcanism is absent between about 28S and 148S


of active volcanism on the depth of the brittle–ductile

transition in the crust, which generally is not

significant at distances larger than 10 km from the

craters (Ito, 1999 and references therein). Further-

more, the lack of correlation between plate contact

resistance and the assumed slab dip (Meijer, 1995;

Meijer et al., 1997) has been considered as unex-

pected, as it is known that areas with a dflat-slabT inthe Andes have an intense crustal seismicity relative

to the activity observed in areas of dnormalT sub-

duction, which suggested that compressive stresses

should be higher in the dflatT portions of this region.The main process assumed to explain flat WBZ is

the subduction of oceanic lithosphere with over-

thickened crust; this process has been assumed in

almost the entire literature related to dflatT subductionbeneath the Andes. Flat WBZ in Ecuador, Peru,

central Chile and Argentina have been related to the

subduction of the Carnegie Ridge, Nazca Ridge and

Juan Fernandez seamount, respectively (e.g., Cross

and Pilger, 1982; Pilger, 1984; Gutscher et al., 1999,

2000b). Within other processes proposed for explain-

ing dflatT WBZ, active overthrusting of the overriding

continent is a main alternative to the effects of

increased buoyancy of the overthickened crust (Vlaar,

1983; van Hunen et al., 2000). Recently, van Hunen et

al. (2002a,b) have provided an accurate analysis of the

effects of these processes in Japan, Peru, central Chile

and Argentina. A significant conclusion is that dflatTsubduction in Peru, Chile and Argentina cannot be

explained solely by subduction of ridges and sea-

mounts, even when high temperatures (N600 8C) areassumed in the models for the basalt-to-eclogite

transition in the subducted oceanic lithosphere. The

extent of the dflatT slab segment in the direction of

subduction could be explained by invoking active

overthrusting of the South American continent, but, as

van Hunen et al. (2002a) underline, the contribution

of this effect is uncertain, at least because it depends

on the hot spot reference frame which validity is

widely debated (Meijer and Wortel, 1992). Also, very

few models describe a dre-steepeningT of the flat slab

after full plateau subduction, necessary to fit several

seismological observations. While this can be due to

the two-dimensional model used, which reduces the

ability of the dflatT slab to separate from the overlying

continental lithosphere, it is also certain that model

results depend largely on this two-dimensionality,

whereas ridges and seamount chains are long and

relatively narrow structures (van Hunen et al., 2002b).

These critical results concerning two major processes

assumed to explain dflatT WBZ in the South American

continent can be taken as an indication of their

unplausible character.

In this work, the geotherms and rheological regimes

of these areas of the Andes are determined on the basis

of heat flow observations and physical properties of

rocks. Values of radiogenic heat generation for the

middle and lower crust, and for the upper mantle, are

.

,

t

t

l

.


those assumed globally. Heat generation in the upper

crust is found under appropriate boundary conditions.

Crustal radiogenic heat generation obtained through

the present analysis is found to be in agreement with

global average ranges measured in old terranes, in

correspondence to the generally old geologic age (e.g.,

Myers, 1975; Ramos et al., 1984; Litherland et al.,

1994; Astini et al., 1995; Casquet et al., 2001) of the

terranes that form the main tectonic structures of these

areas (Figs. 3–5). The potentially seismogenic zones in

Fig. 5. Map of central Chile and Argentina (after Astini et al., 1995) show

The Precordillera, the Cuyania terrane and all the allochthonous Lauren

Cambrian time, and following drift in the Iapetus ocean collided with Gond

accretion of Chilenia in the Palaeozoic. Grenville basement outcrops are e

Palo to the area of Mendoza. The dflatT surface of seimicity at about 100

mainly beneath the geologically old terranes of western Argentina below

between about 288S and 338S.

the crust and continental mantle are described for

different thermal parameters and thickness of the crust.

Seismicity of the lower crust is thoroughly explained,

and it is shown that seismic activity defining dflatTWBZ is a misinterpretation of what has been seen in

modern seismotectonics as a rare kind of seismicity,

that is, the genesis of earthquakes in the subcontinental

uppermost mantle without participation of subducted

oceanic lithosphere. Recent seismological studies are

discussed to confirm these results.

ing the allochthonous terranes and the observed or inferred sutures.

tian terranes of western Argentina were rifted from Laurentia in

wana in the Ordovician. The postcollisional interval ended with the

xposed along the Precordillera eastern border from the Sierra Pie de

km depth—observed for about 200 km in the E–W direction—lies

zones of enhanced crustal seismicity. Active volcanism is absent


2. Geothermal features and partition of the

rheological regimes

In the plate tectonics model, the interaction of a

subducting plate with the continental lithosphere is

clearly manifested in the thermal structure of the

system. Subduction of cool oceanic lithosphere cools

the overlying continental lithosphere mainly at dis-

tances less than about 200–300 km from the trench

when dip of the slab has dnormalT values (e.g., Toksozet al., 1971; Honda, 1985; Peacock et al., 1994;

Peacock, 1996; Peacock and Wang, 1999; Davies,

1999). When subduction is subhorizontal, the thermal

structure of the entire continental lithosphere from the

arc to the back-arc may also be affected, and a

temperature decrease reaching maximum values of

about 100–400 8C can be expected in the continental

crust and mantle depending on the model parametri-

zation and thermal processes considered in the

interaction (Gutscher et al., 2000a,b). The effect of a

cool oceanic lithosphere should be greater in the case

of an initially warm continental lithosphere, causing a

significant downward migration of the isotherms. This

is not the case of the areas studied in this work, where

a cold continental lithosphere is assumed to be related

to low crustal radiogenic heat generation.

The unperturbed thermal structure of the continen-

tal lithosphere depends mainly on the surface heat

flow, the radiogenic heat generation and the thermal

conductivity of the crust and uppermost mantle. In

this work, geotherms are computed under steady state

and one-dimensional assumptions. The rationale for

these assumptions will be presented later as remarks

on the results obtained. These geotherms give

approximate knowledge of the thermal structure when

the detailed conformation of the system is unknown.

In regions of rather high heat flow, as the Altiplano–

Puna Plateau, transient thermal effects are strong, and

description of the thermal evolution needs several

variables and processes to be considered (e.g., Le

Pichon et al., 1997).

Surface radiogenic heat generation data are avail-

able in few areas of the Andes. In southern Chile, at

latitudes between 338S and 348S, it ranges from

1.45F0.74 to 2.09F0.43 AW m�3 (meanFS.D.), in

Palaeozoic and Tertiary granitoids, respectively; at

latitudes between 398S and 418S, it ranges from

1.07F0.34 AW m�3 in the western area (Jurassic) to

2.09F0.59 AW m�3 in the eastern area (Miocene)

close to the volcanic chain. A maximum value for the

average crustal heat generation of the Central Andes

between 218S and 278S is 1.3 AW m�3 (Lucassen

et al., 2001).

A variable thermal structure with different parti-

tion of rheological regimes constitutes significant

constraints for the genesis of seismic activity in the

continental lithosphere (e.g., Sibson, 1982; Chen and

Molnar, 1983; Ranalli, 1991; Deverchere et al.,

2001). By considering the thermal and gravity

structures together with parameters derived from

seismological studies and creep experiments, it is

possible to construct rheological profiles of the

lithosphere describing approximately the potentially

seismogenic zones in the crust and upper mantle

(e.g., Fadaie and Ranalli, 1990; Kaikkonen et al.,

2000; Deverchere et al., 2001; Pasquale et al., 2001;

Tejero and Ruiz, 2002). The approach followed here

is the comparison between a linear frictional fracture

criterion describing a brittle regime (Anderson–

Sibson formulation) and a nonlinear flow stress

(the power-law equation for dislocation creep)

corresponding to a ductile rheology (e.g., Chen and

Molnar, 1983; Ranalli, 1991). The fracture criterion

can be set as rza.avz(1�k), where r is the

difference between maximum and minimum princi-

pal stresses, g is the acceleration of gravity taken as

9.78 m s�2, z is the depth, .av is the average density

of the overlying rocks at depth z, and k is the pore

fluid pressure. a is a numerical factor that depends

on the mechanism of faulting. For a cohesionless

material with a coefficient of friction equal to 0.75,

and intermediate principal stress equal to the average

of maximum and minimum principal stresses, atakes the values of 3.0, 1.2 and 0.75 for thrust,

strike-slip and normal faulting, respectively (Sibson,

1974). In the ductile regime, the principal stress

difference is r=(C�1de/dt)1/nexp(E/nRT(z)), where

de/dt is the strain rate, R the universal gas constant,

C the dislocation creep parameter, E the activation

energy for creep, n a stress- and temperature-

independent parameter, and T(z) the absolute temper-

ature at depth z. Creep parameters C, E and n

depend on rock type. The transition from the brittle

to the ductile regime, describing maximum depths

Z( q) of the potentially seismogenic zones in the

lithosphere measured from the surface, is obtained


by comparing the brittle and the ductile critical stress

difference as a function of depth z. The geotherm is

calculated by assuming the general heat flow pattern

(Hamza and Munoz, 1996). When this procedure has

been applied in areas of the Andes with normal or

high heat flow, the rheological results (Munoz, 2000)

have been found to be consistent with other geo-

physical observations (e.g., Lezaeta et al., 2000;

Campos et al., 2002).

3. Low heat flow in the Andes region

The low heat flow observed in some areas of the

Andes region (Fig. 6) has usually been explained as

Fig. 6. Automatic contour map of heat flow for latitudes between 08 and 40

excluded from the data set (Hamza and Munoz, 1996). The map was gener

flow measurements in the Andes inside the contoured areas with qb40

conventional (temperature gradient and conductivity) method, and crosse

temperature) method. In Peru, between about 4.58S and 158S, three of th

located in the coastal region. The eastern extension of the area with qb40 m

temperature data east of about 768W giving a mean heat flowFS.D. of 3

anomaly are constrained by conventional heat flow data from sites in the W

sites are located in the coastal region at distances between 130 and 150

constrained by one conventional observation in the Cordillera; this extensio

map for central western South America of Bahlburg and Herve (1997) tha

and northwestern Argentina. In Argentina, two sites are located in the Pre

conventional data from the Central Valley and Cordillera of Chile and by b

of the anomaly is in some agreement with magnetotelluric studies (Munoz

north of 278S in western Argentina, and where at least for about 180 km so

values. Heat flow measurements are lacking in the Cordillera between 278Srheological situation when the assumed heat flow is larger than 40 mW m

an effect of nearly subhorizontal subduction of the

Nazca plate. Instead here it is proposed that in these

areas the low radiogenic heat generation in allochth-

onous terranes and old geologic structures of western

South America (Figs. 3–5) is a main cause of heat

flow less than about 45 mW m�2. The Argentine

Precordillera and the Chilenia terrane were accreted

to the protocontinent between Ordovician and

Devonian time (Ramos et al., 1984; Dalla Salda et

al., 1992; Dalziel et al., 1994; Astini et al., 1995;

1996; Casquet et al., 2001), an extensive Precam-

brian basement with Palaeozoic cover is distin-

guished from northern to southeastern Peru

(Bellido, 1969; Myers, 1975), and processes of

accretion took place in Ecuador between the Late

8S where low-quality geochemical estimates of heat flow have been

ated by kriging, available in the SURFER package. The sites of heat

mW m�2 are shown. Dots denote data obtained by following the

s denote those obtained by means of the MGT (underground mine

ese heat flow sites are in the Western Cordillera and other sites are

W m�2 in northern and central Peru is constrained by 52 bottomhole

9F7 mW m�2; the southern and the southeastern extensions of the

estern Cordillera and the Altiplano. In central Chile, the heat flow

km from the trench and the northern extension of the anomaly is

n of the anomaly is in satisfactory agreement with the revised terrane

t precludes the extension of the Chilenia terrane into northern Chile

cordillera. The southern extension of the anomaly is constrained by

ottomhole temperature data in Argentina. The northeastern extension

et al., 1992) that indicate low electrical resistivity of the lithosphere

uth of this latitude the lithosphere is characterized by high resistivity

and 338S; models 26–28 (Table 2) are representative of the thermo-�2.


Jurassic and the Early Tertiary (e.g., Megard, 1989;

Litherland et al., 1994; Jaillard et al., 1997).

Particularly, the Argentine Precordillera (Fig. 5) is

a terrane derived from Laurentia and comprising a

basement of Grenville age (c1100 Ma), and

probably linked to terranes of the southeastern

United States as it shows stratigraphic similarities

with the Tennessee–Alabama southern Appalachians

(e.g., Astini et al., 1995, 1996; Casquet et al., 2001).

Low heat flow (b40 mW m�2) in the eastern United

States has generally been associated to low radio-

genic heat generation in the basement of Grenville

age or to regional redistribution of heat through

large-scale groundwater flow (Birch et al., 1968;

Blackwell et al., 1991). At present, the most

extensively studied Grenvillian terrane is the Gren-

ville province of the Canadian Shield (e.g., Pinet et

al., 1991; Guillou-Frottier et al., 1995; Jaupart and

Mareschal, 1999) where heat flow shows a rather

random character with values of 41F11 mW m�2

and with an area weighted average of 39 mW m�2

(Jaupart and Mareschal, 1999). The correction for the

Pleistocene climatic effect on heat flow of the

Canadian Shield (2–3 mW m�2) lies within the

range of measurement uncertainty; this effect was

most prominent in the Northern Hemisphere and

should not affect temperatures in the Southern

Hemisphere. The radiogenic heat generation of

principal rock types with surface abundance larger

than 1% varies from 0.1 to 1.8 AW m�3, with an

area weighted average of 0.8 AW m�3 (Jaupart and

Mareschal, 1999).

Heat flow measurements between about 278S and

338S in the Precordillera of NW Argentina and in

central Chile give low values (21.0–41.9 mW m�2),

but are few in number (7) and not well distributed

(Uyeda et al., 1978a,b; Uyeda and Watanabe, 1982;

Hamza and Munoz, 1996). In Peru, from about 4.58Sto 178S, the general heat flow pattern is one of low

heat flow (b40 mW m�2); in the eastern region, heat

flow increases to 40–60 mW m�2. In southern Peru,

the area of low heat flow is interrupted by a belt with

heat flow higher than 40 mW m�2 (Hamza and

Munoz, 1996). Ecuador is almost not explored for

heat flow, and the general pattern corresponds to an

automatic contour map generated by kriging (Fig. 6)

that gives values between 40 and 60 mW m�2 (Hamza

and Munoz, 1996).

As described later, mantle heat flow in areas of

low heat flow in western South America is found to

be about 12–14 mW m�2. These values are

comparable to those obtained in the Grenville

Provinces of the Canadian Shield and in the

Appalachians (Pinet et al., 1991; Guillou-Frottier et

al., 1995; Lenardic et al., 2000). In an outstanding

contribution, Lenardic et al. (2000) have shown that

mantle heat flow appears to be low and spatially

uniform across tectonic provinces ranging in age

from 400 to 2700 Ma, and that surface heat flow

variations in stable continental regions are not due to

mantle heat flow but to variations in crustal heat

generation. It can be suggested that this can also be

applied to Grenvillian and Precambrian terranes of

western South America, although in this case

thermal events have occurred recently in the geo-

logical record and these areas cannot be considered

as purely stable areas. The significance of crustal

radiogenic heat generation and the character of some

thermal events as highly localized phenomena are

apparent in zones with active volcanism. From

magnetotelluric imaging of the lithosphere in the

southern volcanic zone of Chile, high temperature

and presence of fluids can be inferred only close to

the active volcanic chain and main faults, respec-

tively (Munoz et al., 1990; Brasse and Soyer, 2001),

in agreement with the observed surface heat gen-

eration that is generally between 0.7 and 1.5 AWm�3 west of the area with active volcanism (see the

preceding section). This observation agrees with

short-period group-velocity tomography of Rayleigh

and Love waves where none of the active volcanic

centres along the Andes appears to produce distinc-

tive velocity anomalies (Vdovin et al., 1999), and

where the volcanic centres are seen to block the

propagation of shorter-period Lg waves (Rial and

Ritzwoller, 1997), which suggests that the thermal

anomalies associated to volcanism must be small.

Also, it is noted that in zones of active volcanism in

Japan, the cutoff depth of seismicity is observed at

distances very close (about 10 km) from the craters,

beyond which tectonic earthquakes are generated in

the brittle crust (Ito, 1993). Then it is proposed that

the past high heat flow that could be inferred from

volcanism prior to about 10 Ma in old terranes of

Peru, Chile and Argentina was probably a singularity

within broad areas where the regional thermal


structure was largely controlled by low crustal heat

generation.

4. Background for calculations

4.1. Thermal conductivity and computation of

geotherms

In the present work, unlike previous geothermal and

rheological research in the Andes (Munoz, 1994,

2000), where a very weak temperature-dependent

conductivity was assumed in the lower crust—a model

broadly used in geothermal research (e.g., Chapman

and Furlong, 1992; Rudnick et al., 1998)—both

temperature and pressure dependence of thermal

conductivity are considered on the basis of improved

models and experimental observations on crustal rocks

(Sass et al., 1992; Seipold, 1998). In the upper mantle,

thermal and pressure effects are also taken into

account, following a recent model for thermal con-

ductivity in the mantle, based on phonon lifetimes from

infrared reflectivity (Hofmeister, 1999), and which

significantly differs from previous models in that the

radiative conductivity is negligible below 1200 8C.Taking hlc as the depth to the top of the lower crust

and h as the total thickness of the crust, for depth z

[m] and temperature Tb [8C], the thermal conductivity

, [W m�1 8C�1] for crustal rocks (zVh) at temper-

ature Tb and pressure P(z) is: ,(Tb,P)={,(0)/[1.007+Tb(0.0036�0.0072/,(0))]}(1+1.5�10�6 z),

where ,(0)=,(25)[1.007+25(0.0037�0.0074/,(25))]and ,(25) is the conductivity at room conditions (25

8C and atmospheric pressure). For rocks lying at depth

Vhlc, ,(25) is 3.0 or 3.3 W m�1 8C�1; for lower

crustal rocks, ,(25) is 2.6 W m�1 8C�1.

In the mantle, with g=9.78 m s�2, and taking .av

[kg m�3] as the average density of the overlying rocks

at depth z,

, Tb;P� �

¼ 4:44� 0:0044 Tb � 127� ��

� 1þ 0:046.avgz10�9� �

;

for zNh and TbV327 8C;

, Tb;P� �

¼ 3:56� 0:00203 Tb � 327� ��

� 1þ 0:046.avgz10�9� �

;

for zNh and 327 8CbTbV727 8C;

, T b;P� �

¼ 2:75� 0:00088 Tb � 727� ��

� 1þ 0:046.avgz10�9� �

;

for zNh and 727 8CbTbV1227 8C.The geotherm is calculated by an iteration algo-

rithm (Chapman and Furlong, 1992) that follows the

variations of conductivity and heat generation in the

crust and upper mantle. In a layer of constant heat

generation A and constant thermal conductivity ,,temperature Tb and heat flow qb at the bottom of the

layer of thickness Dz are:

Tb ¼ T t þ qtDz=, � ADz2=2,

qb ¼ qt � ADz;

where T t and qt are the temperature and heat flow,

respectively, at the top of the layer. Thermal con-

ductivity effects, ,(Tb,P), are incorporated at each

step in an iterative loop and the computations are

carried out with a 100 m depth increment.

4.2. Pore pressure

Pore pressure is taken to be hydrostatic (k=0.37) inthe upper crust and suprahydrostatic (k=0.90) in the

middle and lower crust. For the mantle, models with

both regimes of pore pressure will be presented. Nur

and Walder (1990, 1992) have suggested the existence

of pore pressure close to lithostatic in the deep crust,

as well as a time dependence of porosity and

permeability involving episodic variation of pore

pressure leading to a pulse of fluid release followed

by a long period of no or little fluid, during which

fluid pressure recovers. Crustal seismic reflectors may

be high pore pressure zones related to subhorizontal

detachment zones (Jones and Nur, 1984; Nur and

Walder, 1990). It should be noted that seismic

reflectors have been determined in the crust of

western Argentina at 308S (Zapata, 1998).

Low-Q factor of seismic wave attenuation in the

crust is also consistent with suprahydrostatic pore

pressures (e.g., Nur and Walder, 1992). Low-Q factor

in the mantle could also indicate the effect of water—

that has to be separated from the effect of temper-

ature—on seismic wave propagation (Karato and

Jung, 1998; Karato, 2000). Low-Q raypaths from


subcrustal earthquakes in central Peru and from

crustal and subcrustal earthquakes in the Sierras

Pampeanas of western Argentina (Chinn et al.,

1980; Whitman et al., 1992, 1996) can be seen as

indicating that a relatively high content of water is

retained in the continental lithosphere of these

regions, thus enhancing anelasticity with respect to

that determined for other areas of the Andes.

4.3. Strain rate

Shortening in the Sierras Pampeanas (Argentina)

during the last 10 Ma (Jordan and Allmendinger,

1986; Jordan, 2000) indicates a strain rate between

about 10�16 and 10�17 s�1. At 308S, Allmendinger et

al. (1990) estimate a crustal shortening corresponding

to a strain rate between 10�15 and 7.6�10�16 s�1,

with the east-verging Precordillera of Argentina

accounting for large part of the total shortening (strain

rate between 5�10�15 and 8.4�10�15 s�1). These

values are presented as over-estimates because of their

assumption regarding the maximum thickness of the

present crust and because any magmatic input was

neglected even though the ages of the volcanic rocks

overlap in time with the age range of at least 30% of

the shortening.

The short-term strain rate based on seismic moment

tensors of earthquakes occurred in the central Andes

between southern Colombia and central Peru—and

comprising areas of normal and low heat flow—is

1.9�10�16 s�1 when the seismogenic layer is taken to

be 40 km thick (Suarez et al., 1983). Seismic moment

tensor solutions have also been obtained by Dewey

and Lamb (1992) splitting earthquakes in the Andes

into two groups—those occurring in the plate-boun-

dary zone that rests directly on the subducted Nazca

plate (fore-arc earthquakes) and continental crustal

earthquakes occurring elsewhere in the plate-boundary

zone above the asthenosphere. For a seismogenic layer

of 50 km thickness (Dewey and Lamb, 1992), the

short-term strain rates between 58N and 358S range

from 1.7�10�16 to 6.5�10�16 s�1 in fore-arc regions,

and from 3�10�17 to 7�10�18 s�1 in continental

regions like the ones examined here. Dewey and Lamb

(1992) noted that short-term moment release can be

consistent with the long-term pattern even though it

may account for a fraction of the deformation required

to accommodate the full plate convergence.

Liu et al. (2000) examined tentatively the defor-

mation across the Andes in relation to a much larger

crustal shortening indicated by GPS data. In this

work, model results mainly for strain rates of 10�17

and 10�16 s�1 will be presented. These are values that

appear currently in the literature concerning the

rheology of the lithosphere and that are consistent

with the former data on deformation of the Andes. For

North America, Thybo et al. (2000) assumed strain

rates to be 10�16 s�1 in regions of normal and high

heat flow, and 10�17 s�1 in regions of low heat flow.

Zoback and Townend (2001)—under a thermal and

rheological parametrization which differs from para-

metrizations used here, and assuming a tectonic

driving force of 3�1012 Nm�1—found that for

surface heat flow of 60F6 mW m�2 and for near-

hydrostatic and suprahydrostatic pore pressure in the

upper and in the lower crust, respectively, strain rates

are likely to be less than 10�17 s�1, whereas they are

approximately equal to 10�15 s�1 under near-litho-

static conditions in the upper crust. For heat flow

lower than 50F5 mW m�2, Zoback and Townend

(2001) obtained strain rates lower than 10�20 s�1

under either pore pressure regime.

5. Thermal and rheological parametrization

In order to describe the partition between brittle

and ductile regimes in these areas of the Andes,

geotherms are computed for heat flow in the range of

35–50 mW m�2 for step-wise distribution of radio-

genic sources in the crust, and, in some cases, for

comparing results, for an exponential distribution in

the upper crust. The thickness of the crust is taken

from gravity studies (Feininger and Seguin, 1983;

Fukao et al., 1989; Introcaso et al., 1992), and three

layers of increasing density with depth are considered

(see Table 1). k is taken to be 0.37 and 0.90, for

hydrostatic and suprahydrostatic pore fluid pressures

in parts of the lithosphere, respectively. As pointed out

previously, temperature and pressure dependence of

conductivity are considered for both the crust and the

mantle. The values of conductivity assumed at room

conditions (Table 1) are consistent with those meas-

ured in these areas and with values assumed globally

for geothermal modelling (e.g., Uyeda et al., 1978a,b;

Hamza, 1982; Pinet et al., 1991; Chapman and

Table 1

Parameters adopted for the calculation of the strength envelopes and

Z( q)

.,kg m�3

,, W m�1

8C�1a

G, AW m�3 {E (kJ mol�1); log

C (MPa�n s�1); n}

Upper

crust

2700 3.0–3.3 see

Table 2

{219; �2.9; 2.4}

Middle

crust

2800 3.0–3.3 0.4–0.6

(0.7)

{230; �3.0; 3.0}

Lower

crust

3000 2.6 0.2–0.4 {268; �2.8; 3.3}

Upper

mantle

3300 see

Section 4.1

0.02 D {535; 4.5; 3.5}

W {498; 2.6; 4.5}

W* {471; 3.3; 4.0}

Temperature at the surface: 0 8C; 10 8C (high; low altitudes)

Pore pressure: 0.37; 0.90 (see Section 4.2)

Strain rate: 10�17 s�1 (10�16 s�1)

(D,W,W*): creep parameters for dry (D) and wet (W,W*) rheologies

assumed for the mantle.a Values of conductivity (,) are at room conditions (,(25) in

Section 4.1).


Furlong, 1992; Munoz and Hamza, 1993; Rudnick et

al., 1998; Jaupart et al., 1998; Kaikkonen et al., 2000).

Studies of radiogenic heat generation are lacking in

the areas here studied, and few measurements should

not enable to have an approximate picture of the

distribution of this parameter as follows from its large

variation in geologically old terranes. Therefore geo-

therms are computed assuming a heat generation (G)

of 0.4–0.6 and 0.2–0.4 AW m�3 for middle and lower

layers of the crust, respectively, and of 0.02 AW m�3

for the upper mantle (Rudnick et al., 1998), and the

radiogenic heat generation of the upper layer (A) is

found from geotherms which give appropriate thermal

gradients in the crust/mantle boundary and at about

100 km depth, whereas solutions that do not satisfy

this condition are discarded. This is done by consid-

ering the curves for upper and lower boundaries of

geothermal gradients in the mantle derived by

Magnitsky (1971). For a nearly thermally unperturbed

uppermost mantle, this procedure is consistent with

the constraint that the conductive geotherm, continued

downwards, must intersect an adiabatic geotherm

(Rudnick et al., 1998). In most cases, this results in

mantle heat flow of about 12–14 mW m�2.

Rheological parameters are taken from different

sources and compilations (Carter and Tsenn, 1987;

Wilks and Carter, 1990; Ranalli, 1991; Mackwell et

al., 1998). A summary of the creep parameters {E (kJ

mol�1); log C (MPa�n s�1); n} used in the calculation

of the strength envelopes is given in Table 1. For the

upper two layers of the crust, the rheological para-

metrization corresponds to rocks of intermediate

composition. Most important is the rheology assumed

for the lower crust and for the upper mantle. The

following rheologies were considered for the lower

crust: Adirondack felsic granulite {243; �2.1; 3.1},

Pikwitonei mafic granulite {445; 4.1; 4.2}, diabase

{268; �2.8; 3.3}. For the upper mantle, dry (D) and

wet (W, W*) rheologies of dunite are used. The strain

rate de/dt is taken to be 10�17 s�1 or 10�16 s�1.

Models of P-wave velocity in the Argentine Precor-

dillera and in central Peru (Smalley et al., 1993;

Dorbath, 1996) indicate that the lower crust is

predominantly felsic, with probably some mafic

granulite in the lowermost crust. In models shown in

Table 2 (see also Fig. 7), the rheology of diabase has

been used because it gives intermediate values for the

maximum depth Z( q) of seismogenic zones in the

lower crust as compared with those obtained with the

rheologies of Adirondack and Pikwitonei granulites.

With log C=�2.8 in the rheology of diabase,

variations of this parameter of F0.5 produce only

slight variations (1–2 km) in Z( q). An intermediate

wet rheology for the upper mantle (W) has been used

in most of the models shown in Table 2; this is

consistent with conclusions about the rheologies of

suboceanic and subcontinental upper mantle and the

rheology of the upper mantle under island arcs (Karato

and Wu, 1993). Other effects of variations in thermal

and rheological parameters are discussed later.

6. Seismogenic zones in the continental lithosphere

6.1. Modelling results and thermo-rheological

variations

Several results of models for different surface heat

flow ( q) and thickness of the crust (h) in areas of the

Andes are shown in Table 2 and Fig. 7. Radiogenic

heat generation in the upper crust is found to be

between 0.7 and 1.2 AW m�3, which considering the

values assumed for the middle and lower crust (0.4–

0.6 and 0.2–0.4 AW m�3, respectively) results in an

average of total crustal heat generation in the range

from 0.43 to 0.73 AW m�3, in agreement with average

Table 2

Thermo-rheological and seismic parameters

Area q, mW m�2 h, km A, AW m�3 TCMB, 8C TBD, 8C (N to T) Z( q), km (N to T)

Ecuador (18N–28S)(1) S.W.P. 40 35 1.0 366 – – – – 35–35

597–532 85–71

(2) E.W.P. 40 40 1.2 409 402–356 39–33

594–530 86–70

(3) E.W.P. 45 40 1.2 508 413–296 31–21

621–556 59–48

(4) S.W.P. 45 50 1.0 527 344–303 27–23

607–548 73–56

Peru (58S–178S)Zone 58S–138S(5) S.W.P. 35 33 1.0 289 – – – – 33–33

578–518 113–96

(6) S.D.H. 35 33 1.0 289 – – – – 33–33

578–532 113–100

(7) S.W.P. 40 33 1.0 363 – to 354 33–32

599–534 77–65

(8) S.W.P. 40 33 1.2 335 – – – – 33–33

593–528 89–75

(9) E.W.P. 35 50 0.7 422 396–292 46–39

590–526 97–79

(10) S.W.P. 35 50 0.7 438 397–347 44–37

591–531 91–75

(11) E.W.P. 40 50 0.9 521 413–298 36–24

606–547 73–57

Zone 158S–178S(12) S.W.P. 40 50 0.8 492 408–295 38–25

600–540 82–64

(13) S.W.P. 40 50 0.7 505 405–359 37–32

604–548 76–61

(14) S.W.P. 40 65 0.7 612 340–291 30–25

616–No T 66–No T

(15) S.W.P. 45 50 1.0 549 346–304 27–23

614–554 66–51

Chile and Western Argentina (27.58S–338S)(16) S.W.P. 35 45 0.7 412 398–346 43–36

590–528 91–75

(17) S.W.P. 35 45 0.7 412 398–346 43–36

616–551 98–81

(18) S.W*.P. 35 45 0.7 412 398–346 43–36

563–509 84–70

(19) S.D.H. 35 45 0.7 412 330–239 34–23

590–540 91–78

(20) S.W.P. 35 45 0.7 393 – to 347 45–38

586–524 100–82

(21) E.W.P. 35 45 0.8 424 404–352 42–35

590–529 91–74

(22) S.W.P. 35 45 0.9 394 – to 347 45–38

585–525 98–81

(23) S.W.P. 38 45 0.8 427 401–350 41–34

589–532 92–75


Table 2 (continued)

Area q, mW m�2 h, km A, AW m�3 TCMB, 8C TBD, 8C (N to T) Z( q), km (N to T)

Chile and Western Argentina (27.58S–338S)(24) S.W.P. 40 45 0.9 443 402–354 39–33

593–535 86–70

(25) S.W*.P. 40 40 0.7 349 – to 343 40–39

552–499 96–81

(26) S.W.P. 45 60 0.9 608 351–298 28–23

618–No T 63–No T

(27) S.W.P. 50 50 1.2 547 349–306 26–22

610–551 68–51

(28) S.W.P. 50 60 1.0 709 350–299 24–20

No N–No T No N–No T

q: surface heat flow; h: thickness of the crust; TCMB: temperature at the crust/mantle boundary; TBD: temperature at the brittle–ductile transition

resulting from the assumed rheology and relevant conditions; Z( q): maximum seismogenic depth (measured from the surface) resulting from the

rheological model and the characteristic geotherm. Mechanisms of faulting: N: Normal; T: Thrust (for strike-slip faulting, intermediate values of

those for N and T are obtained). For TBD and Z( q), the first entry in each model corresponds to the crust; the second entry corresponds to the

mantle. Dashed lines in the TBD column indicate that in the crust there is no transition to the ductile regime for the relevant mechanisms of

faulting. In the TBD column, indications like d� to 354T mean that there is no brittle-ductile transition in the crust for N. dNo N - No TT (or dNoTT) in the TBD and Z(q) columns means that the mantle is in the ductile regime for these mechanisms. In the models, S stands for a step-wise

distribution of radiogenic elements in the crust; E stands for an exponential distribution in the upper crust. For S models, A is the mean

radiogenic heat generation in the upper crust; for E models, A is the surface radiogenic heat generation. W (W*) and D: wet and dry mantle

rheology; H and P: hydrostatic and suprahydrostatic pore pressure in the mantle. In each area section, the range of crustal thickness h describe

characteristic thicknesses of the regions of low heat flow. Results from one section can be applied to other one when the observed or inferred

pair ( q,h) is similar.


ranges for old terranes of South America and North

America presented by Artemieva and Mooney (2001,

Fig. 11). It is to be noted that the averages of

Artemieva and Mooney (2001) are likely to be

representative of the eastern South American con-

tinent, the old Andean terranes not being considered in

their analysis. For suprahydrostatic pore pressure in

portions of the lithosphere, the temperature of tran-

sition from a brittle to a ductile regime in the crust lies

between 340 and 413 8C for normal mechanism of

faulting (N), and between 291 and 359 8C for thrust

faulting (T). In the mantle the transition temperature

lies between 552 and 621 8C, and between 499 and

556 8C, for N and T, respectively. Transition temper-

ature for strike-slip faulting has intermediate values as

compared to those of N and T.

The lithospheric strength, or integrated stress level,

for suprahydrostatic dominating conditions and a wet

rheology for the upper mantle, is generally between

5�1012 and 7�1013 N m�1 for all the models shown

in Table 2. These values are of the same order of

maximum tectonic forces found in plate tectonics

under planetary assumptions (e.g., Turcotte and

Schubert, 1982). It should be noted that in the Andes

the orientation of the regional maximum horizontal

stress does not seem significantly affected by the

change of strike of the mountain belt, and that in areas

of presumed plate contact (i.e., dflatT WBZ), it does

not appear to be different from areas to the east where

the South American plate overlies the asthenosphere

(Assumpcao, 1992). Forces arising from the conver-

gence of the Nazca plate may not be the only major

contributors to the lithospheric stresses of the over-

lying continental plate. Lithospheric heterogeneities

and a heavy, cold lithospheric root can give rise to

large tectonic stresses (Fleitout and Froidevaux, 1982;

Assumpcao, 1992).

In the models of Table 2, layers of some few

kilometers thickness, with increased heat generation

(1.0–1.5 AW m�3), can be superposed in the upper

crust, or replace part of it, with the effect of increasing

Z( q) by some kilometers in the crust and the mantle

when the thermal conductivity of these layers is close

to the thermal conductivity assumed for the upper crust

in the models. Changing the strain rate of 10�17 s�1 by

an order of magnitude, to 10�18 and 10�16 s�1, has the

Table 3

Summary of results for Coulomb–Navier failure criterion, wet

rheology of the upper mantle and a strain rate of 10�17 s�1

A: heat generation in the upper crust corresponding to constrained

models

ZC: Maximum depths to potentially seismogenic zones in the crust

ZM: Maximum depths to potentially seismogenic zones in the

mantle

NOSM: A non seismogenic mantle is possible according to the

models

Central Ecuador

A: 1.0–1.2 AWm�3 ZC: 21–35 km ZM: 70–86 km

Northern and

northern–central

Peru

A: 0.7–1.2 AWm�3 ZC: 32–46 km ZM: 75–113 km

Southern–central Peru

A: 0.7–1.0 AWm�3 ZC: 23–38 km ZM:NOSM–82 km

Central western

Argentina

(Precordillera)

For wet (W and W*)

rheology of the

upper mantle:

A: 0.7–0.9 AWm�3 ZC: 33–45 km ZM: 70–100 km

For dry (D) rheology

of the upper mantle: ZM: 78–91 km

Northern–central Chile

(Cordilleran area)

A: 0.7–1.2 AWm�3 ZC: 20–45 km ZM:NOSM–100 km.

With a strain rate of 10�16 s�1, ZM increases by about 5–7 km.

Other failure mechanisms (see Section 7.3) and broadening of

the thermal parametrization could also increase ZM by several

kilometers.


effect of decreasing and increasing, respectively, Z( q)

in the mantle by about 5–7 km. A summary of results

for a strain rate of 10�17 s�1 and for wet rheology of

the upper mantle is presented in Table 3.

The computed geotherms, particularly for the

Precordillera of Argentina, show similar features to

those obtained for the Grenville region in North

America when several variations of the thermal

parameters—surface heat flow, heat generation, con-

ductivity and mantle heat flow—are considered (e.g.,

Lamontagne and Ranalli, 1996; Lamontagne, 1999).

Temperatures in the crust and in the crust/mantle

boundary (Table 2) are seen to depend mostly in the

surface heat flow, while variations in conductivity and

heat generation cause temperature fluctuations within

about 20–50 8C when surface heat flow is about 35–

45 mW m�2 (see also Jokinen and Kukkonen, 1999).

Following the suggestion that deformation in the

cold, shallow upper mantle occurs by diffusion

rather than by dislocation creep (Karato and Wu,

1993), this was explored for different creep param-

eters and grain size. In diffusion creep, r depends

on grain size (d) as dc, where cc2.5. Both for wet

and dry rheologies in diffusion creep (Karato and

Wu, 1993), no significant effect on the maximum

seismogenic depths in the mantle is obtained for

grain size z10�3 m. For grain size of 10�4 m, Z( q)

decreases by about 10–15 km, and an increased

strain rate or other failure mechanism should be

needed in the mantle for Z( q) reaching values of

90–100 km. The relativization of the lithospheric

pressure–depth relationship (Petrini and Podladchi-

kov, 2000) was also found to be not significant for

deformation observed in these areas. Thickness of

the brittle layers in the lower crust and upper mantle

decreases by only 1–2 km even when pressure

gradients are close to values twice the lithostatic

gradient. Due to the effects of pressure on con-

ductivity, temperature decreases at the crust/mantle

boundary by about 4–15 8C.

Fig. 7. Strength envelopes for Andean areas of low heat flow and different c

each model, the left curve corresponds to normal faulting (extension); the r

the strength envelopes drawn as thick curves indicate potentially seismogen

the ductile regime. CMB is the crust/mantle boundary. In the models show

the upper crust and suprahydrostatic in the middle and lower crust and in th

the lower crust, the rheology of diabase is used, giving thicknesses of the

compared with results obtained using the rheologies of Adirondack and P

6.2. Patterns of observed seismicity and Z( q) values

In areas of low heat flow (40–45 mW m�2) of

Ecuador (18N–28S), constrained models give values

of A between 1.0 and 1.2 AW m�3. Gravity and

seismic studies indicate a thickness of the crust in

these areas of Ecuador between about 50 and 35 km;

southward of Quito, at about 18S, the crust may be

thicker (Feininger and Seguin, 1983). From Z( q)

rustal thickness corresponding to some of the models of Table 2. For

ight curve corresponds to thrust faulting (compression). The parts of

ic zones in the crust and the mantle. Thin curves show the domain of

n, the strain rate is 10�17 s�1, and the pore pressure is hydrostatic in

e mantle. Wet rheologies of dunite are used for the upper mantle. For

potentially seismogenic zones that are of intermediate magnitude as

ikwitonei granulites.


values, it is noted that the entire thin crust (35 km) and

at least the middle crust when the crust thickens to

40–50 km are potentially seismogenic. In Ecuador, for

zones where heat flow is taken to be about 40 mW

m�2, there is strong indication that the continental

upper mantle is seismogenic down to depths of 70–85

km; in zones of higher heat flow, Z( q) decreases to

about 50–70 km. These values of Z( q), both in the

crust and the mantle, agree with recent analysis of the

distribution of seismicity (see caption to Fig. 3) in

areas of Ecuador (Guillier et al., 2001).

Values of A obtained for Peru range from 0.7 to 1.2

AW m�3 for heat flow between 35 and 45 mW m�2.

The morphology of the Peruvian Andes is very

complex (cf. Dalmayrac et al., 1980), and several

characteristic values of crustal thickness h in the

northern and central areas (58S–138S) and in the

southern zone (158S–178S) are considered (see Fukao

et al., 1989). For values of h different to those of the

models in the Peru section of Table 2, models with

appropriate low heat flow in the Chile–Argentina

section may be considered. In the northern and central

areas, in zones of heat flow of about 35–40 mW m�2

and with h not larger than 40 km, most of the crust is

potentially seismogenic, and earthquakes in the

continental mantle can occur down to depths of about

80–110 and 65–100 km, for N and T mechanisms,

respectively. For h larger than 40 km, the lower crust

or at least part of the middle crust can be seismogenic,

and values of Z( q) for the mantle are of about 75–100

and 60–80 km, for N and T, respectively. In the

southern zone (158S–178S), the middle crust can be

seismogenic and Z( q) for the mantle is generally not

larger than 65–85 km. Below in this zone, the ductile

regime dominates. As for Peru, in other areas, when

heat flow is higher than about 45 mW m�2, the

seismogenic zone in the mantle reduces to a minimum

or is absent unless a more dextremeT rheology were

chosen to represent the mantle flow. In Peru, fault

plane solutions indicate both normal and strike-slip

faulting at depths between 90 and 120 km (Stauder,

1975; Hasegawa and Sacks, 1981; Tavera and Buforn,

2001). In a microseismicity study in central Peru

(Suarez et al., 1990), crustal seismicity in the high

Andes is relatively low with the exception of the

activity observed in a terrane of Precambrian gneisses

in the Eastern Cordillera. Focal depths of crustal

earthquakes range generally from 15 to 35 km, and

some occurrences beneath the sub-Andes appear to be

as deep as 40–50 km. Most of the intermediate-depth

earthquakes ranging in depth from 85 to 110 km

(Suarez et al., 1990) are located beneath the zone of

enhanced seismicity in the crust.

Solutions for western Argentina (Precordillera)

between 318S and 328S give values of A between

0.7 and 0.9 AW m�3 when h is 45 km. Models 17–

20 in Table 2 are variations of model 16 (M16). In

this model, the radiogenic heat distribution in the

three layers of the crust and in the upper mantle are

0.7, 0.4, 0.3 and 0.02 AW m�3, respectively. In M17,

the strain rate is 10�16 s�1 when temperature is

larger than 450 8C. In M18 and M19, W* and D

rheologies are used in place of W for the upper

mantle. In M20, a layer of 2 km thickness with

radiogenic heat generation of 1.2 AW m�3 is lying at

the top of the crust, and could represent the effect of

some kind of sediments or of more radiogenic bodies

on the geotherm and the rheological regime. In M21,

M23 and M24, variations of about 0.10–0.15 AWm�3 are included in the lower layers of the crust.

Values of Z( q) in the crust agree with seismological

results in the Precordillera and Sierras Pampeanas of

Argentina between latitudes 318S and 328S (Fig. 5)

where h is about 45 km (Smalley and Isacks, 1990;

Introcaso et al., 1992) and where the maximum

seismogenic zone is defined at 35 km depth (Smalley

et al., 1993) and at 40 km in earlier studies (Smalley

and Isacks, 1990). A large part of the seismicity at

100 km depth—observed for about 200 km in the E–

W direction—lies beneath geologically old terranes

of western Argentina below the area of enhanced

crustal seismicity. At this depth, the dominant focal

mechanism is strike-slip. Teleseismic data indicate

some activity below the zone at 90–130 km depth,

between about 150 and 180 km, and above this zone

from about 50 km depth (Cahill and Isacks, 1992). A

relocation of earthquakes in this area defines a

seismic zone at about 90–110 km depth, with foci

eastwards from this zone reaching about 190 km

depth (Engdahl et al., 1998). In the relocation

process (Engdahl et al., 1998), almost no seismic

occurrence has been left above and below the zone

at depth of 90–110 km. Z( q) should rapidly

decrease, getting shallower, where heat flow

becomes larger than 45 mW m�2 as follows from

the general heat flow pattern (Fig. 6). This is


consistent with the observation of Smalley et al.

(1993) that north of 318S and south of 328S in this

area, seismic activity in the crust decreases strongly.

In central Chile (27.58S–338S), potentially brittle

mantle could exist in some areas of low heat flow

(b45 mW m�2), while the ductile regime dominates

in zones of increased heat flow or crustal thickness.

The constrained models in the area of anomalous

heat flow in Chile and Argentina yield values of A

between 0.7 and 1.2 AW m�3.

The overall patterns of seismicity in the crust and

in the upper mantle at 75–110 km depth are indicative

of processes acting in the direction proposed by

Gerbault et al. (1999) in which crustal faults

accumulate at the trough of folds, as a continuation

of mantle faults to the surface, and where pre-existing

zones of weakness are not essential for triggering

development of folding.

7. Remarks on the calculation of geotherms and

rheological uncertainties

7.1. Steady state

A steady state thermal regime has been assumed

for calculating the geotherms for different regions of

the Andes. This is in consonance with the idea that

in plate tectonics the thermal regime of the

approaching continental block at the oceanic side

hardly differ from the usual thermal regime in

tectonically undisturbed areas (e.g., Cermak and

Bodri, 1996). The assumption of a thermal steady

state is an approximation that has proven to be

useful in studies of the thermal regime and rheology

of the lithosphere (e.g., Hamza, 1982; Fadaie and

Ranalli, 1990; Cermak et al., 1991; Seno and Saito,

1994; Balling, 1995; Liu and Zoback, 1997; Meiss-

ner and Mooney, 1998; Hyndman and Lewis, 1999;

Zoback and Townend, 2001).

Still, it is worth to emphasize that the validity of

thermal steady state conditions is limited and that it

involves complex processes and other types of steady

states characterizing the dynamic system of tectonics

and denudation. These steady states are related to the

accretionay and the erosional fluxes, the evolution of

topography and the spatial pattern of exhumation

(Willett and Brandon, 2002). Chapman and Furlong

(1992) have explored the thermal effects of several

tectonic processes on steady state geotherms and

showed that these effects are very demonstrative of

the large deviations from equilibrium geotherms in the

lower crust during even relatively simple tectonic

processes. These processes are difficult to be quanti-

fied in most areas, but at least an approximation

regarding the thermal effects of uplift and denudation

in some regions of western South America has been

provided by Henry (1981). For sites at high altitudes

(4–5 km), the uplift and denudation effects on surface

heat flow results in a reduction of about 12–20% in

southern Peru and of about 12–17% between latitudes

278S and 348S; for lower altitudes, the correction for

uplift and denudation amounts to a few percent

(Henry, 1981). This means that surface heat flow

values as large as 40–45 mW m�2 in some areas of the

Andes could also be consistent with seismic activity in

the lower crust and at 100 km depth.

7.2. One-dimensional temperature modelling

Temperatures in the crust and upper mantle were

calculated using a one-dimensional model, with

thermal parameters that vary only in the vertical

direction. It is known that differences in temperatures

with respect to results obtained using two- or three-

dimensional models are generally due to different

distribution of radiogenic heat generation and to the

variation of conductivity chosen for constructing the

different models (Balling, 1995; Hyndman and

Lewis, 1999). For a geotraverse in Northern Europe,

comparison between 1-D and 2-D temperature

modelling indicates that depth differences between

the corresponding isotherms are generally in the

range of 1–4 km, with some maximum peaks of

about 9 km, and with the isotherms obtained by

interpolating 1-D results commonly lying deeper

than isotherms obtained by 2-D modelling (Baumann

and Rybach, 1991). From these results, the max-

imum decrease of Z( q) in the mantle could be of

about 7 km if 2-D temperature models were

considered; in fact, Z( q) decreases not more than

1–2 km for common depth differences (1–4 km)

between isotherms that are based upon 1-D and 2-D

models. In the crust, the effect of 1-D and 2-D

temperature modelling on Z( q) is generally not

significant.


Though the calculation of the geotherm is

subjected to several conditions and procedures,

Kaikkonen et al. (2000) have shown that the strength

envelopes resulting from standard calculation of 1-D

geotherms as the steady state temperature solu-

tions—as done throughout the present work—are

very similar to those based on the Monte Carlo 2-D

steady state simulation of the geotherms (Jokinen

and Kukkonen, 2000), where uncertainties in geo-

thermal parameters are taken into account by

changing their values randomly within their natural

variation. For regions of low heat flow, differences

between the strength envelopes are almost not

significant, with the Monte Carlo simulation of the

geotherms giving slightly smaller strength values in

the ductile crustal domain and slightly larger values

in the mantle. These differences do not import any

significant variation in Z( q).

7.3. Failure criterion

The Coulomb–Navier frictional criterion has been

applied at any depth where the material is not ductile

although it is experimentally confirmed only to

pressures corresponding to middle crustal depths

(Byerlee, 1967; Jaeger and Cook, 1979; Ranalli,

1997). The failure criterion with a static coefficient

of 0.75 implies high stress differences at middle to

lower crustal depths (e.g., Lamontagne and Ranalli,

1996), but there are indications that this coefficient

decreases with increasing pressures (Jaeger and Cook,

1979). Plastic yielding (Ord and Hobbs, 1989) and

high-pressure failure in the case of hydrostatic pore

pressure and for thrust and strike-slip faulting

mechanisms (Shimada, 1993) are mechanisms that

could increase the thickness of seismogenic layers in

the lithosphere, and their effect would be to decrease

the pressure dependence of strength as depth

increases. Both mechanisms could then support the

generation of earthquakes in the deep crust or in the

upper mantle even when a higher heat flow were

assumed in the models (Table 2) for these areas of the

Andes.

7.4. Brittle–ductile transition

The term brittle–ductile transition has been

criticized because it can produce confusion between

the mode and mechanism of deformation, and so the

term brittle–plastic transition has been proposed to

denote the entire transition from purely brittle to

purely plastic behaviour, thus encompassing the

semibrittle field where deformation involves both

plastic and brittle mechanisms (Rutter, 1986; Scholz,

1990; Kohlstedt et al., 1995). Although constitutive

equations do not exist for semibrittle behaviour,

Kohlstedt et al. (1995) have found that if the depth

to the bottom of the seismogenic zone is determined

by the transition to the stable frictional sliding regime,

then that depth will be shallower than the depth of the

transition to the purely plastic regime. The related

velocity weakening–strengthening model (Tse and

Rice, 1986; Scholz, 1990; Kohlstedt et al., 1995)

seems more suitable to describe the depth distribution

of earthquakes in some areas (Lamontagne and

Ranalli, 1996; Kaikkonen et al., 2000), although the

brittle–ductile model applied in different tectonic

settings has provided satisfactory results (e.g., Cloe-

tingh and Banda, 1992; Seno and Saito, 1994;

Nyblade and Langston, 1995; Liu and Zoback,

1997; Pasquale et al., 2001; Deverchere et al.,

2001). In most seismic regions, both the brittle–

ductile and the velocity weakening–strengthening

models fit as a first approximation the depth

distribution of earthquakes (Lamontagne and Ranalli,

1996).

8. Seismological evidence for the absence of flat

WBZ

The database of the National Earthquake Informa-

tion Center (NEIC) of the U.S. Geological Survey is a

source that gives some seismological evidence for the

existence of earthquake foci beneath the maximum

seismogenic depths Z( q) in the upper continental

mantle. In Argentina, just beneath the area with a dflat-slabT (Cahill and Isacks, 1992; Smalley et al., 1993)

examined previously, and for the period between

years 1990 and 2001, the NEIC database reports tens

of occurrences at depths between 150 and 220 km,

with some foci reaching 250 km. The seismic activity

is concentrated between 80 and 100 km (number of

occurrences is one order greater than in deeper levels).

The situation is similar but less apparent in central

Chile. The seismic activity below 80–100 km depth,


as reported in the NEIC database, is much lower in

Ecuador and Peru. Usually, it has been concluded that

hypocentres below this depth range and that recorded

above it and reaching the crust/mantle boundary are

poorly constrained in focal depth (Barazangi and

Isacks, 1979; Engdahl et al., 1998). This certainly can

be the case, but these focal depth determinations

should not be ignored in the analysis of the spatial

distribution of seismicity. In fact, the relocation

procedure can be extremely complex in regions like

western South America (Engdahl et al., 1998). In

Peru, the trend of seismicity that includes both the

high-quality and the poorly located hypocentres

(Barazangi and Isacks, 1979) describes much better

what could be the real situation, with most of the crust

and a layer of the upper continental mantle being

seismogenic, and below with a seismogenic zone

poorly defined by few hypocentres between 150 and

300 km depth. In the area of Ecuador where a 80 km

deep flat-slab was proposed (Gutscher et al., 2000a,b),

it has been shown subsequently that the continental

lithosphere is seismogenic down to depths of 75 km

(Guillier et al., 2001). From the projections of

hypocentres onto vertical planes (Guillier et al.,

2001), it is seen that the real WBZ in that area of

Ecuador dips with angles between about 458 in the

north to about 308 in the south (Fig. 2). The real WBZ

can be defined only through intensive local seismo-

logical studies; even so, only few seismic occurrences

are observed between 80 and 150 km depth (Guillier

et al., 2001). The flat-slab described for explaining the

present adakitic arc magmatism of Ecuador (Gutscher

et al., 2000a,b)—see Fig. 2—is an artifact due to

consider only earthquakes recorded worldwide (Guil-

lier et al., 2001). The seismic activity between depths

of 50 and 75 km beneath the Andes of Ecuador

(Guillier et al., 2001) is occurring in the mantle; at

these depths, under any reliable thermal and rheo-

logical regimes, the crust cannot be seismogenic. As

in Ecuador, in other areas studied here, the seismic

activity in the crust is bounded by dipping planes

beneath which a large part of the subcontinental

mantle earthquakes occur. In central western Argen-

tina, large part of the seismic activity in the crust—

recorded locally (Smalley et al., 1993) and also as

reported in the NEIC database—takes place beneath

the eastern side of the Precordillera and beneath the

Pampean range Sierra Pie de Palo. As described

earlier, northwards of the zone of high activity

observed between 318S and 328S there is a decrease

in number of earthquakes. The decreasing activity is

also apparent from the NEIC database. This can be

due to increased heat flow, special fluid pulsation

phenomena (cf., e.g., Nur and Walder, 1992) and to

the geometry of fault systems and sutures. Nucleation

of crustal and upper mantle earthquakes occurs close

to the level of highest strength where distorsional

strain energy is maximum (Bullen and Bolt, 1985).

Several independent seismological observations

are significant as evidence for the absence of flat

WBZ. In Chile (26.58S–28.58S), local seismicity

recorded by a temporary network describes accu-

rately a WBZ that is not subhorizontal as proposed

by previous models, and earthquakes down to 250

km depth are observed (Monfret et al., 2000). A

marginal evidence for the absence of flat WBZ can

be taken from the study of the dslab-pushT mecha-

nisms of earthquakes occurring in central Chile, Peru

and Mexico, where earthquakes are assumed to be

generated inside the subducted plate along the

shallow dipping subduction zone (Lemoine et al.,

2001, 2002). A dslab-pushT mechanism is unusual

near the plate interface, but it is described as a not

rare mechanism in nearly flat WBZ in Chile, Peru

and Mexico. At least in Chile, the dslab-pushToccurrence cannot be explained by the analysis of

the stress transfer during the earthquake cycle

(Lemoine et al., 2001). It is not clear whether the

observed mechanisms are due to occurrences with

focus in the subcontinental mantle, or, as suggested

by Lemoine et al. (2002), are they due to the

heterogeneity of the stress field inside flexed down-

going slabs. In northern and central Peru, for

explaining time residuals observed in seismological

stations, a velocity model for a flat subducting slab

at 80–100 km depth was obtained using a

dfestooningT ray with several reflections inside the

slab, that is, with propagation paths which have long

segments in the colder, higher P-wave velocity

subducting plate (Norabuena et al., 1994). The

model resulted in velocities in excess for the flat

structure that could not be explained by any effects,

and that were found to be inconsistent with studies

of the thermal structure in subduction zones and with

the mineralogy of slabs (Norabuena et al., 1994).

Other models, including P-wave velocity variations


in large volumes of the crust and upper continental

mantle, should be explored as plausible explanation

of the observed time residuals.

9. Conclusions and discussion

Thermal and rheological zonation in areas of

generally low heat flow in Ecuador, Peru, Argentina

and Chile indicate that earthquakes occurring at depths

of about 70–110 km are generated within the subcon-

tinental upper mantle, and thus they are not associated

to flat subduction of the oceanic lithosphere. The

rheological zonation explains the seismic activity in

the deep crust and uppermost mantle, which is a rare

kind of activity in the Earth because of large tectonic

forces required to overcome the cumulative strength of

the lithosphere, these large forces being not available

in other geologically old tectonic settings (e.g., the

eastern regions of North and South America). Then

seismic activity at depths of 70–110 km generated

within the subcontinental mantle may be a singular

characteristic of the Andes and other few regions of the

Earth, while generally this kind of seismic activity is

not observed almost anywhere (e.g., Seno and Saito,

1994; Maggi et al., 2000).

The highest values of maximum seismogenic depth

Z( q) are obtained for upper crustal radiogenic heat

generation in the range from 0.7 to 1.0 AW m�3; these

values are consistent with the nature of geologically

old Andean terranes beneath which the abnormal

seismicity in the lower crust and upper mantle has

generally been observed. Considering the three types

of mechanisms of faulting, the models described show

that the transition to the ductile regime in the

subcontinental upper mantle generally occurs at

temperatures from about 520 to 620 8C, in agreement

with several general studies—e.g., the dmagic tran-

sition temperatureT of 600F50 8C proposed by

Anderson (1995). Several mechanisms which are

difficult to quantify—as high-pressure failure and

plastic yielding—and broadening of the thermal para-

metrization through the addition of relatively thin

layers with heat generation higher than 1.0–1.5 AWm�3, or considering the effect of uplift and denudation,

should imply an increase of Z( q) in several kilometers

beyond the maximum depths obtained throughout the

present analysis. In contrast, excepting the case of

diffusion creep in the upper mantle for grain size less

than 10�3 m, other processes and assumptions would

have the effect of slightly decreasing Z( q).

From the initiation of studies regarding the WBZ in

South America, the segments described as corre-

sponding to dflat-slabT subduction were taken as

present-day analogues of the Laramide orogeny of

western North America that evolved between the Late

Cretaceous and the Eocene (Coney and Reynolds,

1977; Barazangi and Isacks, 1979; Jordan and

Allmendinger, 1986). The formation of the Rocky

Mountains has been studied for a century, but the

controversy about the existence of a dflat-slabT episodeduring the Laramide orogeny has not ended (e.g.,

Bird, 1995). Yet most of the modern research in the

Andes—concerning tectonics, volcanism and the

morphology of the WBZ—has described these seg-

ments in association with Laramide flat subduction

(e.g., Kay and Abbruzzi, 1996; Gutscher et al.,

2000b), although for the Sierras Pampeanas of

Argentina it was recognized that there is not any

regional subsidence as predicted by such phenomenon

(Jordan and Allmendinger, 1986). Also, whereas it

was suggested that the dynamic effects of the

subducted slab on the continent could in part explain

the inconsistence between the crustal thickness

beneath the Sierras Pampeanas with long-wavelength

elevation and with crustal shortening estimates, it has

not been established whether or not the crust

thickened significantly as response to dflat-slabTsubduction (Jordan, 2000).

In areas of the Andes here examined, recent

seismological studies are indicating that the real

WBZ lies below this subcontinental seismogenic zone

and is not subhorizontal. In Ecuador, where until now

the most intensive local seismological observations

for studying this problem have been realized, only few

seismic occurrences are observed below the earth-

quakes with a maximum focal depth of about 75 km

generated within the subcontinental mantle. These few

occurrences, which are localized in planes dipping

30–458, are defining the real WBZ beneath this area

(Guillier et al., 2001), and thus it can be proposed—

considering the database for earthquakes and other

seismological results—that a similar situation should

be found in the low heat flow areas of Peru, Chile and

Argentina if intensive studies using an appropriate

network are performed there.


On the basis of these results, the present-day

analogues of the dflat-slabT episode during the

Laramide orogeny vanish, as vanishes the question

of continental lithospheric loss in association with flat

subduction in central western Argentina (Kay and

Abbruzzi, 1996). As for the active magmatic arc of

Ecuador, new ways have to be explored to study the

evolution and cessation of volcanism in areas of Peru,

Chile and Argentina, and advanced geophysical

research is needed to determine the degree of

annihilation of the real WBZ beneath these areas.

Studies of the spatial distribution of shallow- and

intermediate-depth hypocentres beneath Peru have

been controversial for a long time. When reviewing

the WBZ determinations and the models of orogenesis

for the central Andes, K. Aki noted that a definitive

conclusion about low-angle subduction had to be

postponed, and that bthe questions why the segmenta-

tion occurred and why they subduct with different

angles remain mysteriousQ (Miyashiro et al., 1982).

Though the question of mountain building is multiple

and complex, the non-existence of flat Wadati–Benioff

zones opens the mind to new adventures framing the

volcanic processes and the geodynamics of the Andes.

Acknowledgements

I thank Valiya Hamza and Trevor Lewis for their

collaboration in heat flow research of the Andes.

Antal Adam, Vladimir Cermak, Ilmo Kukkonen,

Robert Pankhurst, Ladislaus Rybach and Seiya Uyeda

for support. Giorgio Ranalli and an anonymous

referee for reviewing the manuscript, and Kevin

Furlong for his comments on the paper.

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