T HERMAL S TATE O F C AMPI F LEGREI C ALDERA
I NFERRED F ROM S EISMIC A TTENUATION T OMOGRAPHY
Salvatore de Lorenzo ( 1 ) , Paolo Gasparini ( 2 ) ,Francesco Mongelli ( 1 ) and Aldo Zollo ( 2 )
(1)
(2)
Dip art i me n to d i Geo lo gi a e Geo f i sic a, U ni ve rs ità d i B ar i, I ta l y
Dip art i me n to d i Sc ie n ze Fi si c he, U ni v ers it à d i Nap o li “Fe d eri co II” , It al y
ABSTRACT
A three-dimensional Qp image of the Campi Flegrei caldera between 0 and 3 km
of depth has been inferred by the inversion of P rise time and pulse width data of
eight y-seven local earthquakes recorded during the last bradiseismic crises by a
local array deployed in the area by the Universit y of Wisconsin.
The availabilit y both of thermal measurements in 5 deep boreholes and of a heat
flow surface map of the area allowed us the calibratio n of a local temperature T
vs. Qp relationship.
The comparison of Qp, Vp and Vp/Vs images combined with hydrogeological
and geochemical data from deep boreholes admitted us to distinguish low -Q
anomalies fluid -dominated by a deep low -Qp anomal y for which t he temperature
effect can be considered dominant. The latter anomal y is located in the same zone
where several authors retain that the volcanic and magmatic activit y migrated after
the Neapolitan Yellow Tuff eruption. Moreover this anomal y includes the are a
where the existence of a magma chamber at depth between 4 and 5 km was inferred
by a seismic active experiment.
INTRODUCTION
Campi Flegrei is a resurgent caldera located inside the Campanian Plain, a
graben-like structure at the eastern margin of th e Tyrrhenian Sea.
The caldera formation is believed to be related to two main explosive
eruptions. The oldest one occurred 37 -39 ky ago (Civetta et al.,1997 ;De Vivo et
al., 1999) and erupted about 200 km 3 of dense rock equivalent (Wohletz et
al.,1999), producing a huge ignimbrite formation (Campanian ignimbrite or
Campanian grey tuff). The youngest one, dated at about 12 ky, erupted about 40
km 3 of dense rock equivalent producing a thick tuff formation (Neapolitan yellow
tuff) (Orsi et al.,1996).
The most recent eruptive activit y is mainl y phreatomagmatic. It is confined
within the most recent caldera and the main eruptive episodes are clustered at 9.5 12, 8-9 and 4-5 ky ago. The last eruption occurred in 1538 and formed an about
150 m high spatter cone (M t. Nuovo).
Archeological and historical evidences indicate that the caldera bottom has
been slowl y sinking in the last 2,000 y. The sinking trend has been interrupted at
least 40 y before the Mt. Nuovo eruption when historical documents clearl y
indicate th at an inflation was underway. The total uplift is estimated 5 -8 m and it
reached a climax shortl y before or at the beginning of the eruption (see Dvorak
and Mastrolorenzo, 1991; Dvorak and Gasparini, 1991) . Ground surface started to
sink again after the er uption. Quantitative measurements of ground sinking began
in 19th Century, when Niccolini(1829 ) measured periodicall y the depth of the floor
of a Roman market (Serapeo) covered by seawater. The first topographic leveling
survey was carried out in 1905. Fur ther leveling surveys were carried out in 1919,
1922, 1953 and 1968. All these data indicated a rather continuous sinking with an
average rate of 1.3 cm/ y ( Berrino et al.,1984 ). Since 1970 two uplift episodes
inverted the previous trend: the first one occu rred from 1970 to 1972, the second
one from 1982 to 1984. The maximum uplift from 1970 to 1972 was of about 1.50
m and was followed by a sinking recovering about 14% of the uplift. The 1982 -84
episode had a maximum uplift of about 1.80 cm, followed by a si nking which up to
now recovered more than 30% of the uplift. The inflating and deflating areas
appears to have a circular simmetry of about 6 km radius with no significant
differences of shape and size between inflating and deflating episodes ( Avallone
et al., 1999). Low magnitude seismic activit y (Ml = 1 -4) was moderate during the
1970-72 inflation episode and very intense during the 1982 -84 episode. The latter
was concentrated at depths shallower than 3 km ( Aster et al., 1992 ) in the zone of
higher uplift gradient (Berrino and Gasparini, 1995 ). The possible occurrence of a
magma reservoir at about 4 km of depth is indicated by the identification of a P SV converted phase in a seismic profile ( Ferrucci et al., 1992 ), by the maximum
depth of earthquakes ( Aster et al., 1992; De Natale and Pingue, 1993 ) and by
petrological and geochemical data (see f.i. Armienti et al., 1983; Civetta et al.,
1991).
The first genetic models of the ground deformation pattern ascribed it to a
pressure increase at the top of a sha llow magma chamber embedded in an elastic
medium (Yokoyama, 1970; Berrino et al., 1984; Bonasia et al., 1984; Bianchi et
al., 1984-87; Dragoni and Magnanensi,1989 ; Dvorak and Berrino, 1991, De
Natale and Pingue,1993; ). DInSAR images collected in the perio d 1993 to 1996
have a very dense space coverage and allow a reliable estimate of the Mogi -t ype
pressure source location. It has been located about 800 m offshore SW of the town
of Pozzuoli at a depth of about 2,700 m ( Avallone et al.,1999 ). Bonafede et
al.(1986) considered the pressure source embedded in a viscoelastic medium,
whereas Como and Lembo (1992) developed a numerical model to simulate the
effects of fracturation and conductive thermal propagation on the ground
deformation. All these models requir e overpressures of several hundreds MPa to
justify about 3 m of maximum ground uplift in an elastic medium. If a viscoelastic
medium was assumed, an unrealistic viscosit y as low as 10 1 6 Pa.s was necessary in
order to keep the overpressure within reasonable values (Troise,1999). Finall y, the
sudden deflation of the area without the occurrence of any eruption cannot be
easil y explained with mechanical source models.
Oliveri del Castillo and his coworkers were the first to propose that the fast
inflation and deflation of the Campi Flegrei caldera can be an effect of variation
of pore pressure related to the temperature of fluids in permeable rocks ( Casertano
et al., 1976; Oliveri and Quagliariello, 1969 ). Bonafede (1991) and De Natale et
al.(1991) showed in more details how the heating of groundwater permeating
pervious rocks can produce the overpressures needed to produce the observed
ground deformation. 1D and 2D numerical models based on the application of non
equilibrium thermodynamics to the interaction be tween a thermal source and the
circulation of groundwater in a permeable system ( Ascolese et al., 1993 a,b, Gaeta
et al., 1997) have shown that hydrothermal systems have the effect of amplifying
the ground deformation. In fact they transfer at shallower de pth and on larger
surfaces the increase of pressure occurring at the base of the layer. Therefore the
motion of the front of overpressure should cause a progressive transfer to
shallower depths of the inflation center. However, this phenomenon has not been
detected by any of the geophysical measurement in the area. De Natale et
al.(1997) argue that this is due to the caldera walls acting as rigid boundaries
which make the shape of the surface deformation area rather insensitive to the
depth of overpressure source. Recently Wohletz et al, (1999) and Orsi et
al.,(1999), performed finite elements numerical simulations accounting for
conductive and convective heat transfer and a variet y of sizes of magma
reservoirs. They conclude that the contribution of convect ive heat transfer is
needed to justify the temperature variations with detph observed in several
geothermal wells.
In order to validate thermal and mechanical numerical models it is necessary to
have a complete overall picture of the thermal regime actual l y existing in the
upper 4 km of the Campi Flegrei caldera and the elastic/anelastic rock properties.
The present models are tested against the temperature profiles measured in deep
geothermal wells drilled in two sites inside the caldera (S.Vito and Mofet e) (Agip,
1987). The measured temperatures indicate that the inner part of the caldera is
actuall y the site of significant thermal anomal y (a temperature of 420°C was
measured at 3 km of depth at the S.Vito 1 well). Consistentl y, surface geothermal
gradients of about 150-200 °C/km are measured in the AGIP deep wells and in
shallow boreholes inside the caldera ( Corrado et al., 1998 ).
The study of attenuation of seismic waves can integrate these spott y and
shallow data and provide an overall picture of the th ermal regime in the upper 3
km of crust inside the Campi Flegrei caldera. In fact, laboratory measurements in
the seismic frequency range ( Kampfmann and Berckhemer , 1985) indicate an
exponential Arrhenius -law t ype increase of intrinsic attenuation with the
temperature of the rocks. Several recent researches in geothermal areas ( Sanders
and Nixon, 1995; Sanders et al., 1995; Zucca et al., 1994; Wu and Lees , 1996)
have shown that the imaging of the qualit y factor Q is an usefull tool to define
location and ex tension of magma bodies in volcanic areas and to detail the fluid
contribution through fractured systems. Usuall y high -degree of cracking or high
effective porosit y ( Zucca et al., 1994; Sanders et al., 1995) and the presence of
water/gas in the fractures l ower both P and S wave qualit y factors Qp and Qs.
At a global scale, a high degree of correlation is found between high heat -flow
and low-Q zones (Sato and Sacks , 1989; Mitchell, 1995; Romanowicz, 1995),
despite of the uncertainties in comparing estimates of Q from different
methodologies (Romanowicz, 1998) and in discerning the contributions of effective
porosit y and temperature ( Liu et al.,1994 ).
The availabilit y of a high qualit y set of 3 components digital recordings
obtained during the 1982 -84 seismic crisis allows to improve our knowledge of the
thermal state of the caldera by performing a tomographic inversion of the intrinsic
attenuation of P -waves. The existence of deep geothermal wells (up to 3 km of
depth) enables us to constrain the obtained th ermal model with in situ measured
temperatures.
The aim of this paper is to present the results of the 3D Qp imaging, to discuss
the local Qp -T conversion and to infer the thermal properties of the deep part of
the caldera.
DATA
The data considered in t his study represent a high -qualit y subset of 87 events
belonging to a microearthquakes set recorded from January till June 1984 by a
local seismic network consisting of 12, three -components, digital (12 bit)
seismographs, installed in the framework of a co operative project between
Universit y of Wisconsin and Osservatorio Vesuviano ( Aster et al.,1992). On each
P direct phase we measured the rise time  (the time lag between the arrival time
of the signal and the first zero crossing time, i.e. half a period) and pulse width T
(the time lag between the arrival time of the signal and the second zero crossing
time, i.e. a complete period). The number of data for each of the selected events
ranges from a minimum of 8 to a maximum of 18, depending on the number of
recordings and on data qualit y. The ray coverage is shown in Fig.1. The total
number of pulse data for our anal ysis is around 1000 (fig.2). The 3D P -wave
velocit y structure inferred by Aster and Meyer (1988) was used to compute travel
times. The short dis tances traveled by the waves (from 3 to 10 km) justify the
assumption of a straight line ray-propagation.
THE INVERSION METHOD
Source parameters and qualit y factor were computed using the method
proposed by Zollo and de Lorenzo (1999). It is a version of the classical “rise time
method” (Gladwin and Stacey, 1974) modified to account for the effects of
finiteness of seismic source and the consequent directivity often observed on
waveforms generated by earthquakes.
The method is based on the numerical solu tion of a set of non linear equations
relating data (rise time and pulse width) to source (radius, dip and strike) and
attenuation (3D distribution of Qp). The fault is modeled as a Sato and Hirasawa
(1973)(S&H) circular crack, and the attenuating medium i s discretized in
rectangular blocks.
The inverse problem can be stated in the following way:
- given N microearthquakes recorded at a network of M receivers, the data space
consists of the measurements of rise -times i, j and pulse widths Ti, j (i=1,..N;
j=1,..M i , M i being the number of receivers that recorded the i -th event).
- concerning the source parameters space, for the anal yzed i -th event we need to
estimate the radius  0,i of the S&H circular fa ult (we assume a fixed, known value
for rupture velocit y) and the angles i ,i (respectivel y the dip and the strike of
the fault) defining the spatial orientation of the fault plane.
The whole path attenuation term T/Q is obtained by ad ding up the
contribution T i /Q i of each block crossed by the P -wave ray:
T P Ti

Q i 1 Qi
The source and attenuation model space consists of 3N+P parameters. The non
linear s ystem of equations to be solved is therefore given by:
P T
0,i 0,i
i, j, m

sini , i    i, j 
  i, j
vR
c
Q
1 i  N
m
m 1
P T
1  j  Mi


i, j, m
Ti, j  0,i  0,i sini , i   3.9 
vR
c
m 1 Q m
i , j 
0,
sin   2
c

i, j 0 , c,   1 0,i sin   2
c
 i, j 0,i , c,   1
where:
(1)
(2)
(3)
1  6.601  0.015
 2  0.946  0.015
1  0.051  0.015
 2  0.005  0.015
(4)
where Ti, j,m represents the travel time of the first P -wave generated by the i-th
event recorded at j -th station passing through the m -th block, and  i,i is given
by :
2
2
 Ri , j  1  Ti , j 
 i ,  i   arccos 

2 Ri , j




The search for the best fit model parameters is performed through a two -steps
procedure (fixed source parameters, inver sion of Qp; fixed Qp structure, inversion
of source parameters) based on the search of the minimum of the cost function :
N M


   Wi , j i, jobs  i, jest
i 1 j1
2  WT Ti, j
i, j
obs
 Ti, jest
2 
(5)
where i, j and Ti, j
represent respectivel y rise time and pul se width observed
obs
obs
on the j-th seismogram for the i -th event, i, j
and Ti, j
the corresponding
est
est
theoretical estimation of rise time and pulse width arising from a choice of model
parameters and W i , j and WTi , j the accuracy of the observed quantities. The
exploration of the parameters space is realized throughout the “Simplex Downhill
method” (Nelder and Mead, 1965; Press et al.,1993 ) with initial completel y
random (source or Qp) model.
The choice of the weighting factors depends on the quality of the data. We
attributed a weight of 0.8 to 1.0 to data with signal to noise ratio equal or greater
than 8 and a weights from 0.5. to 0.3 to data with lower signal to noise ratio. In all
the cases a smaller weight was assigned to pulse widths relative to rise -times
because of the different accuracy of the two estimates from waveform data. In fact,
pulse widths are usuall y more sensitive to reading errors than rise times, because
the late part of the pulse is often contaminated from secondary arrivals due to
focusing effects or to P -coda waves. Based on this we decided to fix:
WTi , j = 0.5 W i , j .
INVERSION OF CAMPI FLEGREI MICROEARTHQUAKE DATA:
P-WAVE ATTENUATION I MAGES AND THEIR RESOLUTION
In this paper we deal with the computed attenuation parameters while the
results of the inversion concerning the estimate of source parameters are discussed
by de Lorenzo et al.(1999) . Although the P-wave travel distances are r ather short,
both rise time and pulse width data show a dependence on travel time, even if with
a large scatter [fig.2]. At the microearthquake scale both shape and frequency
content of seismograms can be affected by source directivit y ( Zollo and de
Lorenzo, 1999).
The application of the classical rise time method ( Wu and Lees,1996 ) under the
assumption of homogenous Qp structure produces average residuals of 0.053 s for
rise time and of 0.121 s for the total pulse width. The residuals as a function of
the travel time are shown in Fig. 2.
The application of Zollo and de Lorenzo (1999)’ method under the same
assumption produces a reduction of about 100% of the variance of rise time and
pulse width residuals with respect to the case of punctual and impulsive source.
Interestingl y the apparent Qp values for each source -receiver configuration
(i.e. the best fit value for each considered event) range from Qp=18 to Qp=862.
This extreme Qp variabilit y per event is a clear indication of the high
heterogeneit y of th e attenuation structure. The homogeneous Qp value, obtained by
fixing the source parameters at those previousl y retrieved and inverting the whole
data set is Qp=280  70.
In order to estimate the uncertaint y on Qp the classical approach to the study
of the errors in non linear problems was followed ( Vasco and Johnson,1998). It
consists in mapping random deviations of data in the model parameter space. The
final value of the variance was 1.30 s 2 (corresponding to a r.m.s. =0.037 s). The
present 3D model provides a global variance reduction of 38% with respect to the
previous homogeneous Qp model. (Fig.2)
A classical way to establish the resolution level of three dimensional imaging
consists of establishing the way in which a fixed synthe tic 3D image is recovered
after the inversion of a synthetic data set. A detailed anal ysis of Qp resolution of
the actual data set was performed by de Lorenzo (1999).
In this study onl y the most significant resolution test are described, in order to
give a clear representation of highl y resolved volumes of the investigated medium.
The checkerboard resolution test consists of imaging a synthetic Qp structure
formed by an alternation of low (Q=50) and high Q (Q=600) blocks (Fig. 3). This
test indicates that a higher resolution is expected in the central part of the caldera
at middle depths (1 -2 km). At shallow depths (less than 1 km) a highl y resolved
image is obtained for the whole region with the exception of the peripheral blocks.
At large depths (2 - 3 km) the image is less resolved and only to the central region
appear rather well constrained.
The loss of resolution with depth is a combined effect of the increasing P -wave
travel distances and of the smaller ray sampling for deep blocks. Based on these
tests we infer that the western part of the caldera at depths greater than 2 km is not
resolved by the considered data set.
Main features of the obtained 3D Qp maps
Different fundamental features can be observed from the tomographic 3D Qp
images (Fig.4):
 The shallow layer (from 0 to 1 km depth, in the following referred as layer 1) is
characterized by a central N -S elongated high Qp anomal y which separates two
low Qp areas, one centered NE of the town of Pozzuoli, the other in the Solfatara Agnano area;
 The same trend with smoother lateral variations of Qp occur in the intermediate
layer (between 1 and 2 km depth, in the following referred as layer 2) with
secondary, extremel y localized low -Qp values.
 In the layer between 2 and 3 km depth (layer 3) the tren d changes. An East West trending low Qp anomal y, centered in the central -eastern part of the caldera,
is the most prominent feature.
The comparison of Qp tomographic images with seismic wave velocit y maps
inferred by first arrival time inversion [ Aster & Meyer, 1988] shows interesting
correlations (Fig.4). In particular the low -Qp region in layer 3 matches well both
in location and geometry the low -Vs region. Vp to Vs ratio is about 1.9 at 3 km
depth, and it increases to 2.2 in layer 1. The correlation bet ween Qp and Vs
appears weaker moving toward shallower depths, although high Vs are located on
the N edge of the high Qp anomal y in layers 1 and 2. Aster & Meyer (1988)
attributed the high Vp/Vs value at shallow depths to the effect of fluid -filled
cracked rock materials. The decrease of Vp/Vs with depth and the correlation
between low-Qp and low-Vs in layer 3 suggest that thermal effects are dominant in
this region.
Comparison With Previous Attenuation Measurements in the Area
Previous Q determinations a t the Campi Flegrei were based on the modeling of
the spectral properties of coda waves and direct S waves. It was shown that coda Q
(Qc) is almost constant vs. frequency ( Del Pezzo et al.,1987a; and Aster et
al.,1992). This effect was interpreted as an ev idence that intrinsic attenuation is
dominant over scattering attenuation. This seems to be a characteristics of
volcanic areas where intrinsic attenuation should be strongl y affected by the high
temperature fields produced by melting or intruded hot bodi es.
De Natale et al . (1987) determined an average value of direct S wave qualit y
factor in the area (Qs=110 50) by using a spectral technique. Comparing this
result with the spatiall y averaged Qp obtained in the present study (Qp= 280 70)
we obtain that Qp/Qs=2.6. Regional ( Sato and Sacks, 1989) and laboratory
[Kampfmann and Berckemer , 1985] Q determinations coherentl y indicate that
Qp/Qs2.2-2.6 when the energy is primaril y dissipated by shear deformation.
The consistency of our results with previous Q e stimates is shown also from
the comparison with the the kappa parameter of the Anderson and Hough (1984)’
attenuation model estimated at the Campi Flegrei caldera ( De Natale et al.,1987 ):
k=0.015  0.02 s - 1 for SH waves and 0.008  0.0027 s - 1 for P waves. Assuming a
mean P-wave travel distance of 7 km and Vp=3 km/s an equivalent Q p =R/(v k)290
is obtained, consistent with the average Q p obtained from this study (280 70).
INFERENCES ON THE THERMAL STRUCTURE OF THE CAMPI FLEGREI
CALDERA.
The qualit y fact or is actuall y considered as the seismological parameter more
strictl y dependent on the thermal properties of the rocks (e.g. Mitchell,1995; Sato
and Sacks,1989).
Several laboratory measurements indicate an exponential dependence of Q on
the inverse of the absolute temperature (Arrhenius law). As argued by Romanowicz
(1998) the non linear sensitivit y of Q to temperature implies that attenuation
tomography should be able to resolve hot regions (high attenuation) better than
elastic tomography. Kampfmann and Berckemer (1985) determined an exponential
dependence of Q on temperature by laboratory measurements of Q on rock
materials undergoing forced oscillation in the range 0.03 -30 Hz. For rocks with
Qp/Qs in the range 2.2 -2.6 (as we have inferred for Campi F legrei) they found:
T  1160  150log Qp
(6)
Moreover, they observed that, for temperatures lower than 900 °C, the above
equation holds for any materials irrespective of their chemical composition. Wu
and Lees (1996) imaged the temperature fi eld at depth in Coso (California)
geothermal areas by converting Q in T through the eq. (6).
Actuall y, the main problem in recovering T(Qp) is the need to discriminate the
thermal effect from effects due to variations of permeabilit y of fluids filled rocks .
This problem is relevant in volcanic areas characterized by intense geothermal
activit y. Both laboratory and seismological studies (see e.g. Sanders et al., 1995
and references therein; Zucca et al., 1994) have in fact shown that Qp and Qs
strongl y decre ase also with increasing permeabilit y and fluid circulation in
cracked rocks.
Because of this ambiguit y and of the possible dependence of Q on frequency,
Sato and Sacks (1989) firstl y pointed out the need in calibrating T(Qp)
measurements using local ther mal data.
Our knowledge of the thermal state of the Campi Flegrei arises both from
temperature measurements in five deep AGIP boreholes (fig.5) ( AGIP, 1987 ) and
from thermal measurements in thirt y-seven surface boreholes (up to 140 m of
depth) (Corrado et al.,1998) (Fig.6).
The geothermal surface gradient map compiled by Corrado et al. (1998)
provides clear evidences for strong lateral variation in geothermal gradient, due to
the effect of groundwater motion through the surface aquifers. After removal of
thermal disturbances due to the motion of the water, Corrado et al.,(1998)
estimated an average geothermal gradient in the range 150 -200°C/km.
The residual geothermal map, obtained by subtracting the average geothermal
gradient from the unfiltered geotherm al gradient field, enhances small wavelength
thermal gradient highs located at Mofete, north of M.Nuovo and at the Agnano
crater (fig.6). AGIP MF1 and MF2 deep wells are located on the top of the Mofete
anomal y and MF5 is located on its N edge. On the cont rary, the San Vito wells
(SV1 and SV3) are located in an area where no local thermal anomal y is present.
Temperature and hydrothermal mineral zonation vs depth furnish consistent
information on the geothermal system existing at the Campi Flegrei caldera.
A mineral zonation typical of hydrothermal systems is well developed in the
Mofete wells and, although less clear, in the San Vito wells ( Chelini and Sbrana,
1987). Argillitic and illite -chlorite paragenesis are ubiquitous down to a depth of
750 - 900 m (MF1 and 2) , of 1400 m (MF5) in the Mofete area and down to 1400
m (SV3) – 1750 m (SV1) in the San Vito area. They are due to mineral
transformations at temperatures lower than about 250°C. The produced rock has a
very low permeability and it forms the imper vious cap rock of the geothermal
s ystem. At greater depth neogenic minerals (K -feldspar, adularia, albite, quartz)
are dominant. They are formed by precipitation from hot water cooled by the
contact with the host rocks. This zone (calc -aluminum silicate zo ne) defines the
top of the circulating fluid system. The resulting rock is often highl y permeable.
The
temperatures measured in the wells at the depth corresponding to the
transition from the illite -chlorite to the calc -aluminum silicate zone are 250°C at
Mofete and 220 -270°C at San Vito.
The temperature gradients measured and inferred in wells MF2 and MF5 are
much higher than those in boreholes SV1 and SV3, consistentl y with the average
values observed in shallow wells. MF wells have met important aq uifers running
on the argillitic zone (150 -300 m of depth) and within the calc -aluminum silicate
zone (1250 -1600 m of depth) and also at greater depth. No evidence for
widespread aquifers was found at the SV wells, except for a very shallow water
table in the yellow tuff and pyroclastics.
This confirms that the present fluid circulation scarcel y affects the geotherms
of boreholes SV1 and SV3, which can be assumed to represent mostl y a little
disturbed conductive gradient t ypical of the area. These geotherm s show a mean
gradient of 130 -210 °C/km at depths smaller than 1.5 km and of 44 -64 °C/km at
greater depths. This abrupt change may be explained by a change of the thermal
conductivit y, from typical surficial tuffs values (0.85 W/mK) ( Corrado et
al.,1998) to higher values expected for high temperature hydrothermal altered
rocks (2-3 W/mK).
In order to establish a T vs. Qp calibration curve for Campi Flegrei we
considered the values of Qp obtained by tomographic inversion between 1 and 3
km at the blocks near wells SV1 and SV3 and we correlated them to the
temperatures observed at the same depths. This anal ysis provided the following
empirical equation:
T=713 -61 log(Qp)
(7)
The standard deviation of Qp and the spanned temperature ranges are also shown
in Fig. 7, where the K &B relation (eq.6) is shown for comparison. The shaded area
represents the accuracy of Qp. The values of ln(Qp) obtained by relation (7) are
higher than those obtained by K &B equation in the range of values inferred at
Campi Flegrei. This may be due to the effect of the lithostatic pressure which
contributes to the increase of Qp. If this is the case it is possible to evaluate the
rate of change of Qp with pressure ( ). In fact, this effect may be roughly
expressed by:
(8)
Qp  (gh )
If an average densit y of 2300 kg m - 3 is assumed,  results to be 3.0  102 Pa 1 .
The temperature maps at depths between 1 -2 km and 2-3 km as inferred by the
Qp estimates using eq.(7) are reported in Fig. 8. A positive t hermal T(Qp) anomaly
can be seen East of the Mofete area, in agreement with temperatures measured in
wells. However, the details of this anomal y must be considered with caution
because: (1) resolution in this area is poor, as shown in Fig.3, (2) temperatur e
dependence of Qp cannot be separated from the effects due to the important thick
aquifers found deeper than 1 km.
The most significant low Q, high T anomal y at depths greater than 2 km
extends eastward of the town of Pozzuoli, encompassing an area incl uding the
craters of Astroni, Agnano (where a high thermal gradient was measured, see
Fig.6) and Solfatara. Many recent eruptions occurred in this area and present
hydrothermal phenomena (hot waters, fumaroles) are here concentrated.
The location and geometry of this anomal y well correlate with the EW
elongated region of low -Vs observed by local earthquake tomography [ A&M]
(Fig.4 ).
Several evidences suggest the presence of a local magma reservoir at shallow
depths. Bianchi et al., (1987) and Ferrucci et al.(1992) suggested a top of an
hypotetic magma chamber at about 4 km depth on the basis of deformation and
active seismic data . Wholetz et al. (1999) estimated that the chamber had minimal
volumes of 1000 km 3 and 160 km 3 before the Campania ignimbrite and the
Neapolitan yellow tuff eruptions, respectivel y.
CONCLUSIONS
In this paper we have presented the Q attenuation structure arising from the
joint inversion of rise time and pulse width data of direct P waves. The present
thermal state of the caldera was inferred by calibrating Qp vs T using temperature
vs depth curves of shallow and deep boreholes.
The Qp tomographic images clearl y indicate a change of the anomal y pattern
from the shallow/medium (layers 1 and 2) to the deep layer (layer 3). The
shallow/medium depth region is characterized by a nearl y N trending high Q
anomal y which separates areas of rather low Q. The latter are located nearby the
areas of Mofete and Agnano, where hot fluids have been met by deep drilling.
Instead, a major East -West trending, low Q anomal y is the dominant feature
of the deepest layer. This region lies inside the inner caldera where the recent
(from 11 ky B.P. to present) magmatic activit y has been concentrated ( Barberi et
al.,1991; Orsi et al .,1996) and where inten sive hydrothermal activit y is presentl y
observed.
Seismic velocit y estimates of Aster and Meyer(1988) indicate an anomalous
high Vp/Vs ratio (2.2) at 0 -1 km depth in the area where the eastern, low Q
anomal y is located in the present study. This confirms that the dominant
attenuation effect in this layer can be attributed to a fluid saturated volume.
At larger depths P -wave attenuation and shear wave velocity anomalies have
similar shape and extension, whereas the Vp -to-Vs ratio decreases reaching a valu e
of 1.9-2.0 in layer 3. A similar pattern of decreasing values with depth is shown by
porosit y and transmissivit y measured in the central caldera boreholes ( De Vivo et
al.,1989). All these evidences suggest a dominant temperature rather than fluid
filled crack effect on rock rheology for depths higher than about 2 km. Hence
conduction can be considered as the dominant heat transfer mechanism in this
depth range. This thermal anomal y can be related to the effect of elastic property
weakening of rocks induce d by high temperatures which may be originated by a
deeper located heat source, as suggested by other geophysical and geochemical
evidences.
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FIGURE CAPTIONS
Fig. 1.
A) Earthquake epicenters (February-June, 1984) with other g eophysical results
superimposed: 1 areas of Vp/Vs>1.9 at 1 km depth; 2 areas of Vp/Vs < 1.6 at 1 km
depth; 3 40-cm elevation contours for the 1982 -1985 uplift (data from
Osservatorio Vesuviano: Rapporto Sorveglianza 1985); 4 elliptical trace of the
hypothesized ring fault of Aster and Meyer (1988); 5 1 mgal Bouguer gravity
contours. (after Aster et al.,1992)
B) Horizontal ray coverage for the 87 earthquakes considered in this study.
Fig.2.
A) Pulse width and rise time vs. travel time of the anal yzed data set. Also it is
shown the least squares straight line corresponding to the delta -like sources and
homogeneous Q structure assumption. It can be noted the large scattering of both
the data around the best fitting straight line. This may be "a priori" due b oth to
the source effects and/or to the heterogeneit y of the attenuation parameters of the
crust at the Campi Flegrei caldera.
B) Residuals between rise time (and pulse width) data and the corresponding
"data" predicted by the model as a function of the t ravel time. On the left are
shown the residuals corresponding to the assumption of a point -like source and
homogenous Q model, on the right those corresponding to the assumption of a
finite dimension seismic source and the 3D attenuation structure.
Fig.3
Checkerboard resolution test for the three dimensional case. On the left is shown
the true attenuation structure (low Q= 100, high Q=600), on the right the
recovered Q structure.
Fig. 4.
A) Three-dimensional Vs map at the Campi Flegrei caldera (after Aster and
Meyer, 1988)
B) Three-dimensional Vp/Vs map at the Campi Flegrei caldera (after Aster and
Meyer, 1988)
C) Three dimensional Qp attenuation structure of the Campi Flegrei caldera.
Fig. 5.
Temperature measurements as a function of the depth in five AGIP geot hermal
wells.
Fig. 6
High pass filtered map of the surface geothermal gradient at the Campi Flegrei
caldera (after Corrado et al., 1998)
Fig. 7.
A)Comparison between T(Qp) estimated by Kampfmann and Berckemer
relationship [1985] and T(Qp) values inferred by the calibration procedure
described in the text. Also it is shown the geotherm of the SV1 geothermal well.
A) Comparison between the Kampfmann and Berckemer [1985] relationship and
the T=T(Qp) curve inferred by the local calibration.
Fig. 8.
A) Temperature map at depth between 1 and 2 km as inferred by the local Qp -T
conversion
B) Temperature map at depth between 2 and 3 km as inferred by the local Qp -T
conversion.