Physics of the Earth and Planetary Interiors 124 (2001) 81–93
On the features of the geodynamo following reversals or
excursions: by absolute geomagnetic paleointensity data
Avto Goguitchaichvili a,∗ , Pierre Camps b , Jaime Urrutia-Fucugauchi a
a
b
Instituto de Geofisica, Universidad Nacional Autónoma de México, Ciudad Universitaria, 04510 Mexico DF, Mexico
Laboratoire de Geophysique, Tectonique et Sedimentologie, CNRS and University of Montpellier 2, Case 060, 34095 Montpellier, France
Received 5 June 2000; received in revised form 12 October 2000; accepted 10 February 2001
Abstract
We carried out a Thellier paleointensity study of a ∼3.6 million year Pliocene geomagnetic excursion recorded in a lava
flow succession from southern Georgia (lesser Caucasus). Previous paleomagnetic study [Phys. Earth Planet. Int. 96 (1996)
41] revealed that several consecutive lava flows record an intermediate polarity direction at the base of the section followed
by a thick reverse polarity zone. Samples of 71 from 26 flows from both polarity zones were pre-selected for paleointensity
experiments because of their low viscosity index, stable remanent magnetisation and close to reversible continuous thermomagnetic curves. Altogether 54 samples from 21 flows yielded reliable paleointensity estimates. The mean paleointensity of
the intermediate field is 7.8 ± 2.4 ␮T (three flows). The stable polarity paleointensity is higher with a mean 24.2 ± 8.2 ␮T
(15 flows), which corresponds to a mean virtual dipole moment (VDM) of 4.6 ± 1.8 × 1022 Am2 . This value is significantly
lower than the average Pliocene geomagnetic dipole moment and post-intermediate dipole moments recorded in volcanic
sequences at Hawaii (∼4 Ma) and Steens mountain (∼16.2 Ma). However, our results are quite similar to the post-intermediate
field recorded in Iceland during the Gauss–Matuyama reversal. These results suggest that the regime of the geodynamo following reversals or excursions may vary significantly from one case to the next without any apparent systematic features.
© 2001 Elsevier Science B.V. All rights reserved.
Keywords: Paleointensity; Geomagnetic excursion; Geodynamo; Lesser Caucasus; Pliocene
1. Introduction
The short periods (103 –8 × 103 years after Merrill
and McFadden (1999) and ≥3 × 103 years after
Gubbins (1999)) during which the geomagnetic field
changes polarity are of considerable interest in our
understanding of the physical processes in the earth
liquid core that generate the field. Detailed studies of
∗ Corresponding author. Tel.: +52-56-22-42-30;
fax: +52-55-50-24-86.
E-mail address: [email protected]
(A. Goguitchaichvili).
geomagnetic transitions and excursions have also revealed new features concerning possible core-mantle
interactions (Hoffman, 1992). The virtual geomagnetic
poles (VGP) recorded by sedimentary rocks show the
existence of two opposite preferred longitude sectors
(Laj et al., 1991). In contrast, volcanic rocks, which
are generally more reliable field recorders, indicate
a uniform longitudinal distribution of VGPs (Prévot
and Camps, 1993).
Geomagnetic field intensity should be a decisive
parameter to better understand the field behaviour
during and around reversals or excursions (Camps and
Prévot, 1996). Following Glatzmaier et al. (1999) and
0031-9201/01/$ – see front matter © 2001 Elsevier Science B.V. All rights reserved.
PII: S 0 0 3 1 - 9 2 0 1 ( 0 1 ) 0 0 1 9 0 - X
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A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
Coe et al. (2000), paleointensity during and around
transitions is a fundamental constraint on recent numerical models that promise to provide unprecedented
insight into the operation of the geodynamo. However,
reliable absolute paleointensity is generally much
more difficult to obtain than directional data, even
when the most reliable method (Thellier and Thellier,
1959) is used. This is mainly because only volcanic
rocks which satisfy some specific rock-magnetic criteria can be used for absolute intensity determination
(Kosterov and Prévot, 1999). Judging from scant
available data, there is now a general agreement that
the intermediate polarity geomagnetic field is largely
reduced with respect to the stable, normal or reversed
field (Goguitchaichvili et al., 1999; Quidelleur and
Valet, 1996; Chauvin et al., 1990; Prévot et al., 1985).
Bogue and Paul (1993), based on the study of about
4 Ma volcanic sequence in Hawaii, proposed that the
strength of paleofield following reversals may be unusually strong, probably indicating a prolonged period
of some ‘perturbed’ state of the geodynamo following
polarity transitions. A somewhat similar conclusion
was reached by Valet and Maynadier (1993) analysing
relative paleointensity data. In this paper we present
the results of Thellier–Thellier paleointensity determinations from the ∼3.6 Ma (Camps et al., 1996)
geomagnetic excursion recorded in the southern
Georgia volcanic succession. The volcanic sequence
represents a relatively rapid eruption rate and hence a
high temporal resolution. Camps et al. (1996) carried
out a paleointensity study on the lower Akhalkhalaki
units. However, these results should be considered
as preliminary, because (1) no pre-selection of the
samples was made before undertaking paleointensity experiments; (2) no control heatings (so-called
‘pTRM checks’) were performed during the measurements; (3) experiments were carried out in air, which
can produce some magnetic changes at moderate/high
temperatures; (4) it is hard to estimate the reliability
of these data, because no information about directional changes during paleointensity experiments is
available; (5) obtained results are of bad technical
quality — only 12 specimens out of 40 reported, yield
a quality factor (Coe et al., 1978) above 5; (6) data for
12 lava flows are based on only single determinations.
One of the main goals of the present study was
to supersede the paleoinetensity estimates of Camps
et al. (1996), thought to be of poor quality, by more
reliable determinations. In the present study a considerable effort was spent on the strict pre-selection of the
samples, based on new rock-magnetic measurements.
All heatings (including various control heatings to
detect possible magnetic/chemical changes due to the
heating) were made in vacuum better than 10−4 mbar.
Obtained results are of high technical quality and comparable to other paleointensity data recently obtained
on young lava flows. The natural remanent magnetisation (NRM) fractions used for paleointensity
determination range from 28 to 87% and the quality
factors (Coe et al., 1978) vary between 3.7 and 29.8,
being normally greater than 5. Thus, present results
should be considered as more reliable than previously
reported by Camps et al. (1996). Finally, we gathered previously reported detailed post-intermediate
paleofield records (Tanaka et al., 1995; Riisager and
Abrahamsen, 2000; Prévot et al., 1985; Bogue and
Paul, 1993) in order to obtain some constraints for
the functioning of the geodynamo following reversals
or excursions.
2. Geology and paleomagnetism
Alpine, late Miocene to Holocene compression
due to the northern drift of the Arabian plate towards the stable Russian platform is responsible for
the relief of the great Caucasus. At the front of the
Arabian–Eurasian collision, lithospheric thinning due
to E–W extension linked to the opposite lateral expulsion of the Anatolian and Iranian blocks is at the
origin of the volcanic activity in the lesser Caucasus
(Rebai et al., 1993). Most of the volcanic activity is of
Pliocene to Quaternary ages (Milanovski, 1968). The
Akhalkalaki volcanic area is located in the western
part of the south Georgian volcanic province (Fig. 1).
Camps et al. (1996) studied in detail this 250 m thick
volcanic sequence of some 63 lava flows. The studied
site is located at 41◦ 28 37 N latitude and 43◦ 22 51 E
longitude, ∼1 km S. Se of the village of Thoki and
near the road from Aspindza to Akhalkalaki (Fig. 1a).
The present paleointensity study was carried out on
the lower Akhalkalaki sequence (Fig. 1b, sections W
and Y), which overlies the Goderzi Miocene volcanic
tuffs. Based on two concordant Ar–Ar ages, Camps
et al. (1996) proposed an age of 3.60±0.06 Ma, as the
best estimate of the time of emplacement of the lower
A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
83
Fig. 1. (a) Schematic geologic map showing the location of the Thoki site (modified from Camps et al., 1996). (b) Schematic cross-section
showing location of Y and W sections (also see text).
Akhalkalaki sequence. Upper Akhalkalaki sequence
(section X) which were sampled from the bottom of
a paleovalley uphill (Fig. 1b) most probably were
emplaced in Holocene time.
Four directional groups (DGs after Mankinen et al.,
1985) were identified from bottom to the upper part
of the lower Akhalkalaki sequence by Camps et al.
(1996). The sequence starts with intermediate polarity flows (DG1, four flows) with mean VGP latitude
of 6.3◦ and VGP longitude of 349.9◦ , followed by
another seven intermediate flows (DG2, mean VGP
latitude, 14◦ ). Next, 15 consecutive flows define DG3
with mean VGP latitude of −55.1◦ , which we assign
to the already post-intermediate, reversed polarity
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A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
geomagnetic field. In this paper, we formally use a
VGP cut-off angle of 45◦ to separate intermediate and
stable regimes of geomagnetic field (McElhinny and
McFadden, 1997). The sequence is completed by 14
flows (DG4, also see Table 2) with a mean VGP
latitude of −89.3◦ . Because, the initial polarity of
the geomagnetic event recorded in Thoki volcanic
succession is not known, two possibilities may be
invoked: normal–transition–reverse (N–T–R) geomagnetic transition or R–T–R excursion. If first case
is correct, Thoki geomagnetic event should correspond to either C2An·3n–C2An·2r (dated as 3.363
million year in Huestis and Acton (1997) time scale)
or C3n·1n–C2Ar (dated as 4.096 million year).
Because, Thoki lavas are precisely dated at 3.60±0.06
million year (see Camps et al., 1996), it is unlikely
that Thoki corresponds to a normal to reversed
polarity transition. The geomagnetic event recorded at
Thoki may correspond to some unspecified excursion
which occurred during chron 2Ar of the Gilbert epoch
(Camps et al., 1996; Goguitchaichvili et al., 1997).
coercivity components with MDF around 100 mT
(Fig. 2, sample 91C169C) were systematically
found for site 13W. Continuous susceptibility experiments obtained on the same site (Fig. 2) yield
Curie temperatures (Curie temperatures were determined following the method of Prévot et al.
(1983)) compatible to almost pure magnetite. These
properties may be interpreted as a predominance
of single-domain size magnetic grains of Ti-poor
titanomagnetite or to mixture of single-domain
and superparamagnetic grains, which are suitable material for the Thellier paleointensity
determination.
3. Low-field continuous susceptibility measurements
performed in vacuum (using a Bartington susceptibility meter MS2 equipped with furnace) show the
presence of a single ferrimagnetic phase with Curie
temperature compatible with Ti-poor titanomagnetite (Fig. 2). However, the cooling and heating
curves are sometimes not perfectly reversible.
In all we selected for the paleointensity experiments
71 samples, which belong to 26 lava flows having the
above-described magnetic characteristics.
3. Sample selection
Pre-selection of the flows was mainly based on the
demagnetisation of NRM and temperature dependence
of initial magnetic susceptibility. Magnetic characteristics of typical samples selected for Thellier paleointensity measurements are summarised in Fig. 2 and
could be described in the following way.
1. Samples do not present a big capacity for viscous remanence acquisition. Viscosity experiments
(Prévot et al., 1983) provided viscosity indexes
generally less than 5%, which are small enough to
obtain precise measurements of the remanence during the process of thermal demagnetisation (Prévot
et al., 1985).
2. Selected samples carry essentially a single and
stable component magnetisation, observed upon
alternating field treatment (Fig. 2). A generally
minor secondary component, probably of viscous
origin was present, but was easily removed. The
median destructive fields (MDF) range mostly in
the 40–50 mT interval, suggesting the existence of
‘small’ pseudo-single domain grains as remanence
carriers (Dunlop and Özdemir, 1997). Higher
4. Paleointensity determination
Paleointensity experiments were performed using
the Thellier method in its classic form (Thellier and
Thellier, 1959). All heatings were made in a vacuum
better than 10−4 mbar. At each temperature step, the
specimens were heated twice with an applied field:
positive for the first heating, and negative for the second. The temperature settings were established from
earlier studies of the unblocking temperature spectrum of the Thoki lava flows. The measurements were
carried out in two series: for 35 samples, 12 temperature steps (Fig. 3) were distributed between room
temperature and 570◦ C, and the laboratory field was
set to 30 ␮T. For the remaining 36 samples, only six
steps were distributed between room temperature and
550◦ C and laboratory field was set to 25 ␮T (Fig. 4,
samples 6Y-911621E and 4W-91C062B). Control
heatings, commonly referred as pTRM checks, were
performed after every heating step throughout the
whole experiment. All remanences were measured
using a JR5A spinner magnetometer.
A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
85
Fig. 2. Magnetic characteristics of typical samples, selected for Thellier palaeointensity experiments. Left side: orthogonal vector plots
of stepwise alternating field or thermal demagnetisation (stratigraphic coordinates). The numbers refer to peak alternating fields (mT) or
temperatures (◦ C). Right side: susceptibility vs. temperature curves. The arrows indicate the heating and cooling curves.
86
A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
Fig. 3. The representative NRM–TRM plots and associated orthogonal diagrams for the first series with 12 heating steps (also see text).
In NRM–TRM plots open circles denote to the ‘pTRM’ checks; in the orthogonal diagrams we used same notations as in Fig. 2.
A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
87
Fig. 4. The representative NRM–TRM plots and associated orthogonal diagrams for the second series with six heating steps (samples
6Y-91C621E and 4W-91C062B) and example of worst technical quality determination obtained in this study (sample 8Y-91C318F). Same
notations as in Fig. 3 (* 350◦ C heating step is missing for sample 4W-91C062B due to the ‘accident’ occurred during the measurements
using an automatic data acquisition system).
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A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
Table 1
Thoki basalts paleointensity determinationa
Tmin − Tmax
f
g
q
FE ± σ (FE )
FE ± S.D.
FE 5
6
6
200–521
250–520
20–521
0.65
0.68
0.72
0.71
0.65
0.73
10.7
8.4
19.2
9.48 ± 0.4
9.3 ± 0.5
13.1 ± 0.4
10.6 ± 2.1
11.3
91C042B
91C046E
5
7
200–520
250–540
0.61
0.43
0.69
0.61
13.8
4.2
7.0 ± 0.2
7.3 ± 0.5
7.2 ± 0.2
7.0
3.0
91C075b
91C083D
6
5
200–525
350–550
0.67
0.65
0.63
0.7
7.1
9.7
7.8 ± 1.5
3.8 ± 0.1
5.8 ± 2.8
5.5
151.3
3.3
91C131C
91C134D
9
7
150–520
150–450
0.42
0.48
0.83
0.81
5.7
5.3
48.4 ± 3.0
40.7 ± 3.4
44.6 ± 5.4
44.3
−32.8
151.5
2.2
91C159B
91C161E
91C162b
8
10
7
400–570
300–570
300–550
0.68
0.78
0.46
0.82
0.87
0.8
16.2
29.8
9.9
27.7 ± 1.0
23.1 ± 0.5
32.7 ± 1.2
27.8 ± 4.8
26.4
6
−35.5
149.2
3.0
91C167C
91C168C
91C169E
7
6
7
250–520
350–550
450–570
0.73
0.47
0.74
0.79
0.75
0.82
19.1
5.3
13.6
17.7 ± 0.3
18.1 ± 1.2
20.0 ± 0.9
18.6 ± 1.2
18.4
18W
10
−56.3
182.8
3.5
91C225B
91C229D
7
7
250–500
200–500
0.53
0.71
0.74
0.75
11.3
14.5
17.6 ± 0.6
16.8 ± 0.6
17.2 ± 0.6
17.2
DG4
19W
10
−59.9
180.3
3.4
91C234b
5
200–500
0.72
0.72
9.3
40.2 ± 2.2
DG3
3Y
4
−29.3
153.4
2.4
91C600F
91C601F
91C603F
5
6
5
350–550
300–520
350–550
0.71
0.51
0.69
0.69
0.72
0.71
10.5
6.3
8.4
17.2 ± 0.8
22.4 ± 1.3
16.2 ± 1.0
18.6 ± 3.3
18.3
DG3
4Y
5
−30.9
157.2
3.4
91C606E
91C607F
91C608B
6
9
6
200–550
200–540
200–550
0.42
0.51
0.64
0.72
0.86
0.77
6.4
12.2
11.8
15.3 ± 1.0
17.0 ± 0.6
26.8 ± 1.1
19.7 ± 6.2
20.8
DG3
5Y
5
–30.7
157.3
2.7
91C612F
91C613E
91C616F
91C617C
7
5
5
6
300–540
200–520
200–520
450–560
0.45
0.38
0.39
0.53
0.75
0.72
0.7
0.79
13.5
5.4
8.6
18.8
11.6
18.1
13.5
25.4
±
±
±
±
17.2 ± 6.1
18.0
DG3
6Y
4
−30.3
154.2
1.6
91C621F
6
200–550
0.69
0.73
19.8
24.3 ± 0.6
DG4
8Y
5
−61.5
179.8
1.7
91C317F
91C318F
8
6
150–500
150–400
0.56
0.28
0.77
0.79
13.7
3.7
30.5 ± 1.0
28.1 ± 1.6
29.3 ± 1.7
29.8
DG4
9Y
5
−59.1
184.4
3.8
91C321F
91C322B
91C323B
91C324F
5
7
7
5
350–550
350–550
350–550
350–550
0.48
0.56
0.52
0.49
0.7
0.79
0.8
0.7
10.6
7.5
21.1
10.2
23.4
20.9
31.4
25.3
±
±
±
±
0.7
1.2
0.6
0.9
25.2 ± 4.5
26.4
DG4
11Y
4
−62.8
173.9
11.1
91C332B
91C333C
91C335b
6
11
7
20–520
250–570
300–525
0.32
0.87
0.51
0.65
0.84
0.52
6.5
26.9
7.5
14.8 ± 0.5
17.4 ± 0.5
25.3 ± 0.9
19.2 ± 5.2
18.6
DG4
12Y
7
–61.6
179.5
2.6
91C342F
91C343C
91C344F
5
6
5
200–520
400–550
200–520
0.45
0.82
0.39
0.71
0.77
0.74
6.3
17.1
8
24.5 ± 1.3
25.2 ± 0.9
25.4 ± 0.9
25.0 ± 0.5
25.1
DG4
13Y
10
−62.8
177.3
2.8
91C348F
91C465F
6
8
200–550
300–550
0.77
0.68
0.73
0.81
15.1
12.9
22.8 ± 0.9
25.0 ± 1.1
23.9 ± 1.6
24.2
DG4
14Y
7
−59.3
179.8
1.8
91C359D
91C360E
9
5
150–520
200–520
0.44
0.46
0.72
0.69
7.7
4.8
19.3 ± 0.8
12.5 ± 0.8
15.9 ± 4.8
25.8
Inc
Dec
α 95
DG
Flow
Nd
Specimen
DG1
1W
10
52.9
249.1
4.2
91C031B
91C034C
91C038C
DG1
2W
7
50.2
247.0
2.6
DG2
5W
10
50.5
260.6
DG3
10W
9
−28.1
DG3
12W
10
DG3
13W
DG4
N
0.3
0.9
0.4
0.6
40.2
24.3
A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
89
Table 1 (Continued)
DG
Flow
Nd
DG4
15Y
DG4
DG4
α 95
Tmin − Tmax
f
g
q
FE ± σ (FE )
FE ± S.D.
FE 7
5
6
5
300–540
200–520
300–520
200–520
0.61
0.54
0.53
0.58
0.78
0.67
0.75
0.65
15.2
19.6
5.9
10.8
22.1
17.6
21.0
17.2
±
±
±
±
0.7
0.3
1.4
0.6
19.5 ± 2.4
19.9
91C369b
91C371B
91C372b
91C374C
7
6
8
6
300–575
20–520
300–600
20–520
0.75
0.67
0.64
0.66
0.67
0.77
0.75
0.78
10.7
11.5
6.4
11.9
27.7
22.5
27.5
24.1
±
±
±
±
1.3
1.0
2.1
1.0
25.5 ± 2.6
24.4
91C377b
5
200–500
0.40
0.72
13.2
91.3 ± 2.0
Inc
Dec
Specimen
7
−60.3
180.6
2.1
91C361F
91C362F
91C363D
91C364F
16Y
5
−60.0
182.5
2.4
17Y
4
−61.0
180.7
3.5
N
91.3
a
Directional group number (DG) and lava flow number from Camps et al. (1996); Nd is number of samples used for paleodirection
determination; Inc and Dec flow mean magnetic inclination and declination, respectively; α 95 radius of confidence cone about average
direction; N number of points in the T min − T max interval; Tmin and Tmax the minimum and maximum of the temperature range used to
determine paleointensity; f, g and q NRM fraction, gap factor and quality factor, respectively (Coe et al., 1978); FE the paleointensity
estimate for an individual specimen and σ (FE ) its standard error; FE the unweighted average paleointensity of an individual lava flow, the
plus and minus sign corresponding to the standard deviation; FE a weighed average (Prévot et al., 1985).
b We included some determinations (in total six samples marked by stars) from Camps et al. (1996), because they fulfilled the acceptance
criteria, we imposed in this study (also see text).
Paleointensity measurement data are reported on
the classical NRM–TRM diagram (Figs. 3 and 4).
We accepted only determinations: (1) obtained from
at least five NRM–TRM points corresponding to a
NRM fraction larger than about one-third (Table 1);
(2) yielding quality factor (Coe et al., 1978) generally
above 5; (3) with positive ‘pTRM’ checks, i.e. the deviation of ‘pTRM’ checks were less than 15% and (4)
with reasonably linear Zijderveld diagrams obtained
from the paleointensity experiments. For the best quality samples, the linearity was observed up to 550◦ C
(Fig. 3, samples 12W-91C161E and 4Y-91C607F)
and the control heatings were successful. The worst
technical quality determination obtained in this study
belongs to the sample 8Y-91C318F (Fig. 4) with
quality factor of 3.7. However, paleointensity estimate from this sample is very close to the site mean
paleointensity. Thus, we prefer to keep this determination in our data set. A rather typical ‘concave-up’
behaviour is observed for a few samples (Fig. 4, sample 4W-91C062B), which may correspond to some
irreversible variations of coercitive force (Dunlop
and Özdemir, 1997; Kosterov and Prévot, 1999) at
low/moderate temperatures and can be interpreted
as transformation from a single-domain ‘metastable’
state to multidomain that result in large NRM lost
without any correlated partial TRM acquisition during
the subsequent cooling.
Finally, 54 samples from 21 lava flows (in which
three are of intermediate polarity) yielded apparently
reliable absolute intensity determinations. The NRM
fraction f used for paleointensity determination ranges
between 0.28 and 0.87 and the quality factor q (Coe
et al., 1978) varies from 3.7 to 29.8, being normally
greater than 5 (Table 1). These results correspond to
data of good technical quality. Six comparable technical quality individual determinations from Camps
et al. (1996) data set are also incorporated in Table 1.
Three lava flows are presented by single, but high technical quality determination. These samples, which are
shown in italic at Table 1, are not used in calculating
mean paleointensity or virtual dipole moment (VDM).
5. Results and discussion
The extensive paleomagnetic and paleointensity
studies reported in this paper and by Camps et al.
(1996) provide a detailed vectorial description of the
earth’s magnetic field during the ∼3.6 Ma old geomagnetic excursion. The main results obtained in this
study are recapitulated in Fig. 5, which shows the
variation of paleointensity as a function of elevation,
dependence of VDM versus VGP colatitude and local
field paleointensity versus the angular distance from
the axial dipolar direction. The paleointensity from
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Fig. 5. Summary of paleointensity results for Thoki site: (a) absolute intensity against elevation; (b) local field representation (paleointensity
against angular distance from the axial dipole direction) and (c) virtual dipole representation (VDM vs. paleocolatitude).
the intermediate polarity flows is generally low, with
an average of 7.8 ± 2.4 ␮T (three flows) for directions
more than 45◦ away from the axial dipole field. The
intermediate flow mean VDMs range from 1.98 to
1.12 × 1022 Am2 . Post-intermediate paleointensities
corresponding to the two directional groups DG3 and
DG4 (Fig. 5 and Table 2) are significantly higher
(single determinations are not considered in the calculation). For DG3 (six flows), we found a mean
VDM of 5.5 ± 2.3 × 1022 Am2 and for DG4 (nine
flows), the mean VDM is 4.0 ± 1.2 × 1022 . The mean
VDM from both post-transitional directional groups is
4.7±0.4×1022 Am2 , which is almost half the average
value of the Pliocene VDM (8.2±1.2×1022 Am2 after
A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
91
Table 2
Summary of directional and absolute intensity data (selected using the same acceptance criteria as in the present study) for detailed
post-intermediate recordsa
Region
Age (million year)
Georgia
3.6
Oregon
16.2
Site
Nd
Inc
Dec
α 95
k
Nf
VDM
S.D.
Reference
Thoki (DG3)
Thoki (DG4)
Thoki (all)
15
14
29
−30.2
−60.2
−45.7
152
180.2
162.2
0.9
0.9
6.2
214
271
19.9
6
9
15
5.5
4.0
4.7
2.3
1.2
1.9
This study
Steens
28
66.8
357.1
4.5
38.1
4
6.7
2.0
1,2
Kauai
18
32
353.4
5.7
37.4
12
10.8
1.9
3
Hawaii
3.95
Iceland
∼2.43
Reynivallahals
25
−76.8
175.8
3
93.9
8
3.6
1.3
4
Iceland
∼2.11
Hvalfjördur
30
67.7
16.1
5.1
27
5
5.0
3.5
5
Greenland∗
59.4–60.7
Vaigat
51
−68.6
124.4
3.4
34.4
13
7.4
2.9
6
a
Nd is the number of volcanic units used in the mean paleodirection calculation; Inc and Dec site mean magnetic inclination and
declination, respectively; α 95 the radius of confidence cone about average direction; Nf the number of volcanic units used for the mean
paleointensity determination. The data comes from: (1) Prévot et al. (1985); (2) Mankinen et al. (1985); (3) Bogue and Paul (1993); (4)
Tanaka et al. (1995); (5) Goguitchaichvili et al. (1999); (6) Riisager and Abrahamsen (2000).
Goguitchaichvili et al., 1999). It is worth noting that
a considerable variation of absolute intensity was
detected within the same directional group. Somewhat
similar results were found by Prévot et al. (1985),
studying Steens mountain volcanic succession. Similar to Prévot et al. (1985), we believe that the absolute
paleointensity variation of individual flows from the
same DG may be a real geomagnetic phenomena
rather than an experimental noise, indicating that the
intensity of the geomagnetic field varies faster than
its direction.
One of the main objective in this study was to
try to estimate the state of geodynamo by absolute
paleointensity data following intermediate regime of
earth’s magnetic field, i.e. following reversals and
excursions. Gubbins (1999) proposed that during excursion the field may reverse in the liquid outer core,
which has time-scale about 500 years, but not in the
solid inner core. Both transition and excursion of geomagnetic field involve significant departure of magnetic directions from the usual geocentric axial dipole
and dramatic decreases in intensity. The significant
decrease of paleointensity for intermediate polarity
lava flows is confirmed also in our study in broad
agreement with world-wide data. The occurrence of
relatively strong paleointensities following the intermediate state of the geomagnetic field were reported
by several authors (Fig. 6): Bogue and Paul (1993)
studying ∼4 Ma geomagnetic transition in Hawaii;
Prévot et al. (1985) studying ∼16.2 Ma old Steens
mountain reverse to normal transition; Riisager and
Abrahamsen (2000) studying ∼60.9 Ma, C27n–C26r
reversal (Table 2). Similar observations of a strong
post-intermediate field have been made in relative paleointensity studies on loess sequences (Rolph, 1993;
McIntosh et al., 1996) and deep sea sediments (Valet
and Maynadier, 1993). The authors hypothesised that
a strong post-intermediate field is a systematic feature of geomagnetic reversals and excursions. From
20 post-intermediate Steens paleointensity determination, we selected only four determinations (Table 2)
following the acceptance criteria for paleointensity
data reported in this study. Consequently, the mean
VDM is only 6.7 ± 2.9 × 1022 Am2 , which differs
from the unusually high value outlined by Bogue
and Paul (1993). It is hard to access the reliability
of Bogue and Paul’s data, because no information
about the technical quality of individual paleointensity determinations are available. Riisager and
Abrahamsen (2000) provide this information for their
data, which are apparently of good technical quality.
However, these results should be reconfirmed by additional sampling/measurement, because the studied
samples come from the same small blocks. Thus,
geological continuity of the magnetic signal remains
unchecked. Our paleointensity data do not confirm
the presence of unusually high post-intermediate paleointensity. Moreover, two more post-intermediate
92
A. Goguitchaichvili et al. / Physics of the Earth and Planetary Interiors 124 (2001) 81–93
records, Gauss–Matuyama transition studied by
Tanaka et al. (1995) and Reunion–Matuyama studied
by Goguitchaichvili et al. (1999), (see Table 2), both
yielded relatively low average values of the paleofield
strength. These results suggest that the regime of the
geodynamo following reversals or excursions may
vary significantly from one geomagnetic event to the
next without any apparent systematic features.
Acknowledgements
This study was supported by CONACYT (Project
J32727-T) and CNRS–INSU, program intérieur terre
(contribution CNRS–INSU No. 248). The authors
wish to thank two anonymous referees for useful
comments that greatly improved the manuscript.
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