Geophys. J. Int. (1996) 126,314-344
Seismic moment assessment of earthquakes in stable continental
regions-111. New Madrid 181 1-1812, Charleston 1886 and
Lisbon 1755
Arch C. Johnston
Center for Earthquake Research and Information ( C E R I ), The University of Memphis, Memphis, TN 38152, USA
Accepted 1996 February 29. Received 1996 February 29; in original form 1995 February 6
SUMMARY
The sizes of three major or great historical earthquakes are reassessed using the
isoseismal-area-regression tools developed in Parts I and I1 of this study of stable
continental region (SCR) seismicity. The earthquakes are 1811 New Madrid, central
United States, and its following sequence; 1886 Charleston, coastal South Carolina;
and 1755 Lisbon, oceanic intraplate off the continental shelf of Portugal. The analysis
confirms the large size of these events and for the first time places constraints on the
uncertainty of their seismic moment release. Because of the exceptionally low seismicwave attenuation of eastern North America (ENA), a separate North American
regression of seismic moment on isoseismal area was developed. Additionally, the
unknown western extents of the New Madrid isoseismal areas were calibrated with the
patterns of the M 6.3-6.6 1843 and 1895 earthquakes. Application of Part I1 analysis
procedures with these corrections yields New Madrid size estimates, expressed as
moment magnitude, of M 8.1kO.31 for the 1811 December 16, M 7.8k0.33 for the
1812 January 23, and M 8.Ot-0.33 for the 1812 February 7 principal events. The
Charleston earthquake’s magnitude decreases from M 2 7.4 to M 7.3 f0.26 after
compensation for the effect of coastal plain sediments on its inner isoseismals. Intensity
regressions for Lisbon are calibrated against the isoseismal pattern of the nearly
co-located M 7.8 1969 St Vincent earthquake, which in this case increases the predicted
size of Lisbon from M 8.4 to M 8.7f0.39. These size estimates are supported by data
from independent phenomena: extent and severity of liquefaction, the maximum
distance of induced landslides, and for Lisbon, tsunami wave amplitudes. Estimated
source parameters are controlled by crustal or lithospheric temperature, which governs
the depth extent of brittle faulting. Using estimated continental and oceanic geotherms,
viable fault lengths are 30-80 km for Charleston, 120-180 km for 1811 New Madrid,
and 180-280 km for Lisbon for average displacements of 2-4 m, 8-11 m, and 10-14 m,
respectively, and for average static stress and strain drops. At the estimated seismic
moments of this study, the 1811 New Madrid and the 1755 Lisbon events are,
respectively, the largest known SCR and oceanic lithosphere earthquakes.
Key words: earthquake intensity, earthquake-source mechanism, seismic moment.
INTRODUCTION
The North American continent east of the Rocky Mountain
cordillera is one of nine continental-scale stable continental
regions or SCRs (Johnston 1996a, Part I). The purpose of this
paper is to assess the size and source properties of the largest
of its earthquakes, the 1811-1812 New Madrid sequence and
the 1686 Charleston, South Carolina event (Figs la,b). Also
included is an analysis of the oceanic intraplate 1755 Lisbon
earthquake (Fig. lc), which caused more casualties in a
314
stable continental region than is documented for any other
earthquake. The assessment is in terms of the seismic
moment-isoseismal area regression analysis developed in
Johnston 1996b (Part 11) with auxiliary supporting data from
liquefaction and landsliding effects.
The New Madrid sequence of large earthquakes occurred
in the upper Mississippi River Valley, central United States, in
the winter of 1811-12. They were not systematically evaluated
until a century later (Fuller 1912). Drawing from Fuller, as
well as from newspaper accounts, journals, diaries, and letters,
01996 RAS
Seismic moment assessment of earthquakes-Ill
3 15
37"
36"
35"
!
90"
89"
If 2 4 I
current microseismicity (1974-1993)
area of 1811-1812 intense sand blows
and fissures
I
(
alluvium: current and former flood plains
F]
upland areas
Figure 1. Meisoseismal zones of the three study earthquakes. (a) 1811-1812 New Madrid sequence (adapted from Johnston & Schweig 1996).
Tectonic setting: a failed intracontinental rift, last major activity in the Cretaceous. The liquefaction zone (light shading) simplified from Obermeier
( 1989). Seismicity: all earthquakes (magnitude
1.0-5.0) located by regional seismic networks of St Louis University and the universities of
Memphis and Kentucky. Symbols: CR, Crowleys Ridge; BL, Bootheel lineament; NM, New Madrid; RL, Reelfoot Lake; MR, Mississippi River;
KY, Kentucky. ( b ) 1886 Charleston. Tectonic setting: Mesozoic rifted passive margin. The MMI X isoseismal is from Bollinger (1977). The
principal liquefaction zone is shaded [sand-blow craterlet locations from Amick et al. ( 1990)l. The current microseismicity ( M t 4 . 5 )is located by
the University of South Carolina. The tectonic zone (Marple & Talwani 1993) extends north-northeastwards to about 34.2"N. (c) 1755 Lisbon.
Tectonic setting: compressional deforming zone of oceanic lithosphere (Grimison & Chen 1986; Gordon & Stein 1992) between the AzoresGibraltar plate boundary and the Portuguese continental margin. Bathymetry in metres. Onshore faults and dashed seaward extensions are Nf,
Nazark fault; Tf, Tagus fault; Mf, Messejana fault (Moreira 1989); dashed, 'hypothetical' plate boundary from Purdy (1975); mechanism and
aftershock zone of the M 7.8 St Vincent earthquake of 1969 February 28 from Fukao (1973). Bulls-eye symbols are the different Lisbon epicentres
of Ma (Machado 1966), Mi (Milne 1841), Mo (Moreira 1989), and R (Reid 1914).
-
0 1996 RAS, GJI 126, 314-344
316
A. C . Johnston
(b) Charleston, South Carolina
i0.5"W
80.0"
I
33.0'
Atlantic Ocean
32.5"N
+
current microseismicity
--- --0
1886 liquefaction sites
tectonic zone
25 km
I
I
I
( c ) Lisbon
Figure 1. (Continued.)
0 1996 RAS, GJI 126, 314-344
Seismic moment assessment of earthquakes-Ill
Nuttli (1973), Street (1980, 1982), and Street & Nuttli (1984)
showed that the sequence was dominated by three principal
events: 1811 December 16 (-02:15 local), 1812 January 23
(-09:OO local), and 1812 February 7 (-03:45 local). Based
on the extensive meisoseismal (modified Mercalli intensity
MMI 2X) effects, Richter (1958) qualitatively estimated that
all three exceeded Richter magnitude 8, and Davison (1936)
included them in his book on great earthquakes of the world.
Major aftershocks accompanied each of these ‘mainshocks’,
the probable largest following the first event at -08 : 15 local
on 1811 December 16. There were no known foreshocks to
the December 16 event.
The Charleston earthquake of 1886 August 31 (-21:50
local) occurred in the coastal plain of South Carolina, approximately 26 km northwest of the city of Charleston. The primary
source document for its effects is Dutton (1889), which is a
remarkably complete account published after only three years
had elapsed. The felt and damage data are much more complete
for Charleston than for New Madrid, reflecting the rapid
spread of European settlement in the intervening 75 years. The
Charleston mainshock was preceded by at least several felt
foreshocks (Dutton 1889; Nuttli, Bollinger & Herrmann 1986)
and followed by an active aftershock sequence (Seeber &
Armbruster 1987), although, unlike New Madrid, none was of
sufficient size to be destructive.
There were no perceived foreshocks to the 1755 Lisbon
earthquake, but it had several ruinous aftershocks. The local
time of the first shaking on November 1 in Lisbon is variously
reported: 09 : 30 (Kendrick 1957); 09 :40 (Reid 1914; Davison
1936); 09 : 50 (Officer & Page 1993). The first vibrations usually
were characterized as three distinct shocks separated by intervals of about one minute. By all accounts the duration of the
strong shaking was exceptional: even allowing for exaggeration,
Davison estimates 6-7 min based on reports of multiple
independent observers. The tsunami generated is the strongest
documented in the Atlantic Ocean. Casualties totalled
50 000-70 000 in Lisbon alone, but no reliable breakdown
exists as to deaths and damages from the strong shaking,
tsunami, and fire.
As in Parts I and I1 the ‘size’ of an earthquake in this study
is in terms of the scalar seismic moment M , (dynecm),
expressed
as
log(M,),
or
moment
magnitude,
M=(2/3)log(MO)-10.7(Hanks & Kanamori 1979). For all the
study earthquakes, I also examine other source parameters,
such as stress drop, fault dimension and fault slip, but constraining M , is the principal focus. The treatment of the 1755
Lisbon earthquake is more limited than those of New Madrid
and Charleston. Lisbon presents a number of analysis problems
due principally to its offshore location in oceanic lithosphere,
but its great historical import makes the attempt worthwhile.
3 17
present-day distribution of microseismicity. The standard
mb/mLg estimates, from which other published magnitudes
derive, are from Nuttli (1973) for New Madrid, and Nuttli,
Bollinger & Griffiths (1979) for Charleston.
Nuttli (1983) expanded the mb estimates for New Madrid
and Charleston to include surface-wave magnitude M , and
M , by means of empirical source-scaling relationships for midplate earthquakes. Campbell ( 1986), concentrating only on
Charleston, used Nuttli’s mb estimates but felt that the scaling
to M,, and hence to M, was contradictory and derived his
own. Suggested revisions of the Nuttli (1983) midplate source
scaling are given in Nuttli et al. (1992). The M o (or M/M,)
estimates in all these studies derive from intensity-based mb/mLg.
For the New Madrid sequence the mb/mLg range is 7.1-7.4,
and for Charleston it is 6.6-6.9, with mb 6.7 (Nuttli et al. 1986,
1992) to mb 6.8 (Campbell 1986) favoured. In order to obtain
log(M,) or M from these mb/mLg values, a relation between
mb/mLg and M,, such as that derived in Part I, is necessary
but this introduces additional uncertainty. The published M o
estimates are given in Table 1.
No uncertainties or confidence limits were given for any of
the published values. As noted by Campbell (1986), the original
(also the revised) Nuttli midplate scaling produces a relation
between M , and M o that remains at odds with other global
or regional M,-M, relations (see Part I and references cited
therein), in which M , remains approximately equal to M to
about mid-magnitude 8 level. When divergence does occur due
to M , saturation, M is always greater than M,, just the
opposite of the Nuttli scaling. The methodology developed in
Part I1 and applied here obtains M o from isoseismal areas,
avoids the intermediate correlation with short-period magnitude, and therefore provides a more direct estimate of M o as
well as a physical link to the event’s source parameters.
Quantitative estimates of the magnitude or M , of the 1755
Lisbon earthquake are nearly non-existent. Richter ( 1958)
inferred a magnitude of 8 3/4 from the huge seiche limit and
felt area. Other works (e.g. Officer & Page 1993, Machado
1966) commonly quote magnitude 9 without specifying magnitude type. Abe (1979), in the paper that develops tsunami
magnitude M,, provides the only previous Lisbon M , estimate.
(M, is a moment magnitude derived from regression on
tsunami wave amplitudes.) Abe used the tsunami run-up
heights given by Davison (1936) and a calibration from the
small tsunami generated by the 1969 M 7.8 offshore Portugal
(St Vincent) earthquake to obtain an initial Lisbon magnitude
of M, 8.9. He then halved the run-up heights to approximate
more closely tide-gauge data, yielding a corrected Lisbon value
of M, 8.6. Uncertainties were not formally estimated, but Abe
does caution that the Lisbon tsunami run-up data are mostly
near-field, while the method was developed from far-field
Pacific data.
-
PREVIOUS SIZE ESTIMATES
With two exceptions, all previous estimates of the size of the
New Madrid or Charleston earthquakes are based on intensity
attenuation with epicentral distance and its correlation with
short-period magnitudes (equivalent to one-Hertz mb or mLg).
One exception was the first estimate of the M , of the 1811
New Madrid event by Herrmann, Cheng & Nuttli (1978) from
a linear curve fit to eastern North America M , versus isoseismal
area VI data. Another exception (Gomberg 1992,1993) restricted the inferred M , of the three New Madrid events by the
0 1996 RAS, GJI 126, 314-344
ISOSEISMAL A R E A REGRESSION FOR
SEISMIC M O M E N T
The Part 11 SCR regressions
Part I1 of this study of SCR earthquakes develops regression
formulas for recovering seismic moment M o or moment magnitude M from isoseismal areas Ai. These formulas apply to the
New Madrid and Charleston earthquakes, and, with suitable
calibration, to the oceanic Lisbon earthquake as well. Several
318
A . C . Johnston
Table 1. New Madrid and Charleston: previous seismic moment estimates.
16 Dec 1811 New Madrid
23 Jan 1812 New Madrid
07 Feb 1812 New Madrid
01 Sep 1886 Charleston
mb/mLg=7.2,M,=8.5, Mo=4.0 x lOZ7dyne-cm,M 7.7
M o = 3.2 x loz6dyne-cm, M 7.0
M a x 1.0 x loz7dyne-cm, M 7.3
mb/mL,=7.1, M,=8.4, Mo=2.5 x lOZ7dyne-cm,M 7.5
mb/mLs= 7.4, M, = 8.8, M, = 7.9 x loz7dyne-cm, M 7.9
mb/mL,=6.6,M,=7.5, Mo=2.5 x 1OZ6dyne-cm,M 6.9
mb/mL,=6.7, M,=7.7, M0=3.2x 1OZ6dyne-cm,M 7.0
mb/inLg= 6.8, M, = 7.2, Ma= 7.1 x loz6dyne-cm, M 7.2
functional forms were explored for the Part I1 regressions, but
only the F94 (Frankel 1994) format is used here because of its
physical rather than statistical basis. The Part I1 global SCR
F94 regressions are:
SCR: log(Mo)=17.31 +0.959 log(A)+O.O0126&
SCR: log(Mo)=17.59+ 1.020 ~og(~11,)+0.00139&
SCR: log(Mo)= 18.10+0.971 ~0g(~,,)+0.00194&
SCR: log(M,)= 19.83+0.788 l o g ( A v ) + 0 . 0 0 2 6 0 ~
(1)
SCR: 10g(Mo)=20.23 + 1.032 log(AvI)+0.00176&
+
+
SCR: log(M0)=24.05 + 0.440 log(AVIII)
+0.00586&
SCR: log(M01 = 23.22 0.559 log(Av,,) O.OO328&
The prediction uncertainty, oLgM o , which represents the total
one-standard-deviation uncertainty on log(Mo)from any single
Ai regression in eq. ( l ) , has contributions from the formal
standard error of the regression (a;, M o ) and from the scatter
Appendix Table
of the observations about the regression (ores).
A1 of Part I1 provides a complete tabulation of eq. (1)
regression and prediction uncertainties. To combine results
from individual regressions, the best-weighted average Mbest
and its uncertainty oMbcat
from eqs (15)-( 18) of Part I1 are
appropriate:
(2)
For estimating the total uncertainty of Mbest,the variance of
is taken to equal the random error
systematic error (osYs)*
variance ( c T ~ ~ ~(see
, , ) Part
~
TI), yielding
oMfin,l
= J(‘kh,,,,)2
+ (‘.sys)2
$ OMbSSt.
(3)
For simplicity, eqs (2) and (3) are in terms of M; to convert
to log(Mo)multiply oMbcet
or oMsna,
by 1.5.
Eastern North America ENA regressions
SCR earthquakes are relatively rare; this paucity of data was
a principal rationale behind creating a single global SCR data
set. The continental crusts of the seven SCRs with isoseismal
data are more similar to each other than to the crust of active
tectonic regions, yet intra-SCR differences, related to differences
in tectonic evolution of each continent, do exist. In particular,
Mitchell et al. (1993) and Mitchell (1995) have argued that
cratonjc North America exhibits the lowest shear-wave or L,
anelastic attenuation of any continental region because it was
assembled earlier than other continents and experienced no
Nuttli 1983
Herrmann, Cheng & Nuttli 1978
Gomberg 1992, 1993
Nuttli 1983
Nuttli 1983
Nuttli 1983
Nuttli et aI. 1992
Campbell 1986
major tectonic reactivations. If so, the isoseismal areas of SCR
North American earthquakes should be larger than the SCR
average. Examination of Figs 4-8 in Part I1 suggests that this
is true: the majority of the North American data points lie on
the high-A, side of the best-fitting regressions.
North America, with >50 per cent of the Part I1 SCR
isoseismal database, is perhaps the one SCR with sufficient
Mo-Ai data that it can be analysed independently. I carried
out this analysis for F94 regressions only; the results for
Afelt-AV1are shown in Fig. 2 and compiled in Table 2 under
the same format as for the global SCR regressions of Part 11.
These regressions are identified as ENA (Eastern North
America) to distinguish them from the global SCR formulas
of eq. (1). There were insufficient data to obtain stable ENA
regressions for A,,, and AVIIpIn summary form, for comparison
to eq. (1), the ENA F94 regressions are:
6
E N A log(Mo)= 19.67+0.440 10g(Af~,~)+0.00168
ENA: log(Mo)= 19.15+0.664 10g(A,11)+0.00167&
ENA: log(Mo)= 18.53+0.823 log(AIv)+0.O0188
&
(4)
+
+0.00282
ENA: log(M0)= 20.09 + 0.942 lOg(Av1)+ 0.00244 &
ENA: log(Mo)= 20.29 0.574 lOg(A,)
Development of eqs (4) followed Part I1 for outlier identification, data point weighting, and uncertainty analysis, except
that these regressions are based on an average of -50 per
cent fewer observations than the SCR eqs (1). This leads to
larger regression uncertainties (o:, Mo), which are in part
counterbalanced by reduced data scatter ores.Generally, however, total prediction uncertainties (a;, Mo) (Table 2) are higher
for the ENA than for the global SCR regressions because of
the lower number of observations.
The ENA regressions usually yield lower log(Mo)values for
a given A i than do the global SCR regressions, as expected
from the high North American crustal Q. There has been one
other study of eastern North America Mo-Ai relations
(Bollinger, Chapman & Sibol 1993). but only for A,, and A,,,
and for 4.5 5 M 5 7.5. Their linear AvI regression, based on 21
observations, is included in Figs 2 and 3. Agreement with ENA
A,, is good for log(Mo) 24 ( M 2 5.3), but there are significant
differences at smaller M e Since a stable F94 ENA A,,,
regression was not possible, no comparison at A,,[ can be made.
Because the F94 regression format is derived from the
physical principles of seismic-wave propagation, there is justification for extrapolation beyond the regression data, which is
not the case for a purely statistical regression form, for example
the quadratic of Johnston (1994) or the linear of Bollinger,
Chapman & Sibol (1993). However, because of the lack of
0 1996 RAS, G J I 126, 314-344
Seismic moment assessment of earthquakes-Ill
85
s
27
-I
27
0 North American
26
0 outlier
25
24
UJ
0
26 25 24 -
3 19
-
NA-88-1125'
-
23
-
22
21
4.0
6.0
5.0
7.0
29
-
27
-
I
26 25 24 27
NA-90-1019
3.0
4.0
5.0
6.0
-
7.0
2.0
3.0
4.0
5.0
6.0
7.0
28
-
27
h
26
0
5 25
s
24
m
3
23
22
21
21
2.0
3.0
4.0
5.0
6.0
I
4.5
5.0
5.5
6.0
6.5
7.0
Figure 2. Eastern North America (ENA) log(Mo)-log(Ai) regressions using the F94 functional regression form. The ENA Mo-Ai data are from
appendix Table A3(a) in Part 11. The light dash-dot curves are the 95 per cent ( f 2 4 , Mo) confidence limits of the regressions. Data outliers are
labelled with their Part I1 event numbers. For each regression, Table 2 gives the chi-square and residual statistics and representative prediction
uncertainties for both log(M,) and M. For A,, a linear regression on a slightly different data set is included for Comparison. The last panel provides
a comparison between the ENA F94 regressions and the global SCR F94 regressions of Part 11.
data and the large uncertainties at large Ai, the ENA
regressions are not automatically the first choice for use with
large ENA earthquakes like New Madrid and Charleston. In
the following sections I apply both sets of regressions (eqs 1
and 4)to these earthquakes and examine the differences. This
provides a useful check on the internal consistency of the F94
extrapolation to large Ai.
0 1996 RAS, GJI 126, 314-344
Large earthquake-large Ai extrapolation
The SCR F94 regression uncertainties (Part 11, Figs 4-10)
increase rapidly for log(Afe1,to AIlI)27.0(10 million km2); for
log(A,, to A,,)26.0 (1 million km2); and for log(A,,, to
A w , ) 2 5.0 (100 000 km'). As is evident in Fig. 2, this behaviour
is even more pronounced for the ENA regressions because of
320
A. C . Johnston
Table 2. ENA regression parameters, statistics and uncertainty
(a) Ee_Pression
Ai
kO
type of fit
Afelt ENA F94
A111 ENAF94
AIV ENAF94
Av
AVI
19.67
19.15
18.53
20.29
ENA F94
ENA F94
20.09
kl
k2
v
0.440 0.00168
0.664
0.00167
0.823 0.00188
0.574 0.00282
0.942 0.00244
2
2
xy
P(xJ
Ores
No. of
( ~ o g ~ ~ u n i t outliers
s)
z
41
12
1.088
1.071
0.320
0.382
0.436
0.370
2
1
2.53
2.13
0.247
0.167
0.670
0.565
0.584
1
0
0
2.46
2.54
2.33
33
1.168
42
1.214
22
0.845
0.478
Key: F94 format: log(Mo)= k0 + kl*logAj+ kZ*(Ai)1/2
V = no. degrees of freedom = N - K ,N = no. of data points; K = no. of regression parameters
Ore- standard deviation of residuals; z outlier limit = ZOres
(b) Uncertainty
M
Log(&)
type of prediction
uncertainty
Afelt
AIII
AIV
AV
AVI
3.5
21.30
OM
OlogMo
f 0.341
f 0.512
f 0.464
f 0.696
f 0.431
f 0.647
f 0.511
f 0.767
f 0.510
f 0.766
4.0
22.05
OM
~lOgM0
f0.297
f0.446
f0.308
fO.461
f 0.392
f 0.588
f 0.418
f 0.627
f 0.405
f 0.607
4.5
22.80
OM
OlOgMO
f0.294
f 0.440
rt0.259
f 0.389
5 0.382
f 0.574
f 0.395
f 0.592
It 0.344
f 0.516
5.0
23.55
OM
f 0.294
f 0.441
f 0.262
f 0.393
f 0.383
f 0.574
f 0.395
f 0.593
f 0.331
f 0.497
5.5
24.30
OM
OlogMo
f 0.295
f 0.443
f 0.265
f 0.398
f 0.383
f 0.575
f 0.399
f 0.599
f 0.337
f 0.506
6.0
25.05
OM
=10gMo
f0.300
f0.450
f0.267
f 0.400
f 0.386
f 0.579
f 0.421
f 0.632
f 0.339
f 0.508
6.5
25.80
OM
f 0.311
f 0.466
f 0.278
f 0.417
f 0.398
f 0.597
f 0.476
f 0.715
f 0.359
f 0.538
7.0
26.55
OM
f 0.329
f 0.494
+0.307
f 0.461
rt 0.423
f 0.634
f 0.566
f 0.848
f 0.449
0.674
7.5
27.30
OM
9ogMo
f 0.356
f 0.535
f 0.359
f 0.539
f 0.464
f 0.696
f 0.695
f 1.042
f 0.624
f 0.935
8.0
28.05
OM
? 0.391
a10g.m
f 0.586
f 0.431
f 0.646
f 0.520
f 0.780
f 0.817
f 1.255
f 0.857
f 1.285
8.5
28.80
OM
Ol0g.m
f 0.432
f 0.648
f 0.517
f 0.776
f 0.590
It 0.885
f 0.965
f 1.447
f 1.134
f 1.701
Ologluo
9ogMo
~logluo
fewer control data points in the above ranges. Since the New
Madrid, Charleston and Lisbon earthquakes have isoseismal
areas usually in these ranges, it is instructive to compare their
Ai data with the data of other earthquakes with known M o
before performing a detailed regression and uncertainty assessment. This is done in Fig. 3 with the usual ordinate and
abscissa axes reversed to ease comparison in terms of Ai,which
in this format are represented by horizontal lines when M o is
unknown. The known M o data points are the SCR border
events used in Part I1 (Sd, Sunda Arc; Bh, Bihar; SJ, San Juan)
or instrumental SCR M > 7 events with M determined from
M o (GB, Grand Banks) or from M , (Nn, Nan’ao).
Several general observations can be drawn from Fig. 3 prior
to formal regression analysis. Both Lisbon and New Madrid
clearly separate from Charleston at all intensity levels. Lisbon’s
Afelb AI,, and A,, exceed those of the M 8.3 Sunda event on
the Australian SCR. The New Madrid Afe,, (Table3), too
poorly determined to include, would exceed that of the 1934
M 8.1 Bihar earthquake on the India SCR. At A,, the New
Madrid data straddle Bihar’s, and its A,, exceeds that of this
+
great Himalayan earthquake. Except for Afel, Charleston’s
soseismals equal or exceed those of the low-M 7 Grand Banks,
Newfoundland, Nan’ao, China and San Juan, Argentina events.
The Lisbon A,, is anomalously small compared to its outer
soseismals, probably because a large percentage of it is active
.ectonic crust. Overall, except for the Charleston (and New
Madrid) felt area, these qualitative comparisons are consistent
aith the Lisbon and New Madrid earthquakes having moment
nagnitudes greater than 8 and Charleston having one greater
.han 7. The regression analysis of each event (see below)
substantiates this preliminary conclusion.
THE 1811-1812 N E W M A D R I D SEQUENCE
1Data sources
The intensity distribution of the New Madrid sequence is
xincipally from Nuttli (1973), drawing on Fuller ( 1912) and
iewspaper reports, and from Street & Nuttli (1984), drawing
m Street (1980, 1982, 1984). I evaluated three isoseismal maps
0 1996 RAS, G J I 126, 314-344
Seismic moment assessment of earthquakes -111
M 7.0
7
--
7.5
8.0
.
8.5
9.0
4
~
j
6.8
7'2
321
-
Lisbon A,
-
__ SCR F94
-----
ENA F94
-
+
-
1
6.2
26.0
6.0
26.0
27.0
28.0
29.0
30.0
ENA F94
-----
27.0
28.0
29.0
30.0
Log[M,(dyne-cm)]
Log[M,(dyne-cm)]
-SCR F94
26.0
27.0
28.0
29.0
Log[M,(d yne-cm)]
30.0
Figure 3. Large earthquake extrapolations of the Part I1 SCR and this study's ENA log(M,)-log(A,) regressions. The ordinate and abscissa axes
of Fig. 2 have been reversed to emphasize comparison of data by isoseismal areas. Added to each panel are the SCR border event data points of
Part I1 or other M > 7 instrumental SCR events. Data point labels: Part 11, Table A3(a), GB (Grand Banks, NA-29-1118, M 7.2); Table A4(a), Nn
(Nan'ao, CH-18-0213, M 7.3); Table A4(b), Sd (Sunda Arc, AU-77-0819, M 8.3), Bh (Bihar, IN-34-0115, M Kl), SJ (San Juan, SA-77-1123, M
7.4). Error bars are la uncertainties. Heavy horizontal lines represent the log(A,) for the three study events: if solid, confidence in the value is
indicated; dashes indicate decreased confidence. (a) Extrapolations and data for Afclt (solid symbols) and A,, (open symbols). The Af,,, of Stover &
Coffman (1993) for New Madrid is not included (see Table 3). (b) Extrapolation and data for AIY.(c) Extrapolation and data for Av,. In (b) and
(c) the New Madrid Ai for both ENA (full) and SCR (reduced) regressions (see text) are shown.
based on Nuttli's work for the first and best-documented
1811 December 16 event. [Data in Street (1980, 1982, 1984)
and Street & Nuttli (1984) are not contoured for isoseismal
areas.] The first was the original Nuttli (1973) map, later
modified using additional intensity observations (Nuttli 1981).
Finally, Stover & Coffman's (1993) map, without data points,
differs insignificantly from that of Nuttli (1981), except that a
partial felt limit is added based on (an estimated) two data
points. The Stover & Coffman map is superposed in Fig. 4 on
0 1996 RAS, GJI 126, 314-344
a crustal coda-& contour map (Singh & Herrmann 1983), and
the Nuttli (1981) map is shown in Fig. 5(a). The Nuttli (1973,
1981) generalized isoseismal contours are consistent with Singh
& Herrmann's findings-a much higher anelastic attenuation
to the south of New Madrid than to the north and northeast.
The only other known isoseismal mapping of the 1811-1812
sequence (Stearns & Wilson 1973) used an author-modified
version of the modified Mercalli scale and is not analysed
here.
322
A . C . Johnston
Table 3. Isoseismal area regression: input data.
Event
Afelt
New Madrid: NM1.16 Dec. 1811 i-0215 local)
Nuttli 1973
Nuttli 1981
Stover & Coffman 1993
AIII
AIV
AV
Eastern Areas
~2,124,000 1,447,000
2,126,000 1,667,000
(3,043,000) __ 2,232,500 1,517,000
~~
AVI
AVII
AVIII
735,900
914,000
804,400
299,800
375,600
334,500
69,352
83,300
86,400
Adopted log(Ai) value
full doubling to the west
(6.78)
(6.78)
6.64
6.48
6.21
5.83
5.23
with western reduction factorQ
(6.62)
(6.62)
6.47
6.31
6.10
5.72
5.12
New Madrid NM2.23 Tan. 1812 f -0900 local)
- 2,619,000 1,779,000 1,175,000 563,300 131,300 Street & Nuttli 1984 (estimated, this study)
log(Ai): full doubling to the west
6.72
6.55
6.37
6.05
5.42
log(Ai): with western reduction factorb
6.65
6.49
6.31
5.87
5.21
New Madrid: NM3.07 Feb. 1812 (-03:45 local)
Street & Nuttli 1984 (estimated, this study)
- - 2,129,000 1,210,000 557,500 139,200 log(Ai): full doubling to the west
6.63
6.38
6.05
5.45
log(Ai): with western reduction factorb
6.56
6.32
5.87
5.24
[
1811
:1
Street & Nuttli 1984 (estimated, this study)
- 1,904,000 1,110,000 565,200 238,900
54,100
log(Ai): full doubling to the west
6.58
6.35
6.05
5.68
5.03
log(Ai): with western reduction factorQ
6.41
6.18
5.89
5.57
4.92
Charleston: 31 AUP. 1886 (-21:50 local)
Bollinger 1977 (full doubling offshore)
3,824,000
- 2,817,000 1,540,000 538,600 294,500
91,700
6.58
6.45
6.19
5.73
5.47
4.96
WAi)
Bollinger 1993 (Coastal Plain sediment correction) - - (293,200) 156,460
46,110
__
log(Ai)
(5.47)
5.19
4.66
Lisbon: 01 Nov. 1755 (-09:40 local)
This study
14,830,000 8,100,000 4,680,000 2,040,000
990,600
400,000
173,000
log(Ai)
7.17
6.91
6.67
6.31
6.00
5.60
5.24
St. Vincent: 28 Feb, 1969 (02:40:33 UT)
L6pez Arroyo & Udias 1972
- 2,640,000 1,805,000 1,362,000 1,004,000 275,400 log(Ai)
6.42
6.26
6.13
6.04
5.44
( ) Ai values analyzed but not used for Computing Mfinal
'reduction factors from the 1843 Marked Tree, Arkansas earthquake'sAi
breductionfactors from the 1895 Charleston,Missouri earthquake's Ai
.
-
Isoseismal areas measured from the above sources are given
in Table 3. There were no usable data points to the west due
to lack of settlement, so only roughly one-half of the total
isoseismal areas were constrained by intensity data. Lack of
azimuthal coverage is a common problem for historical earthquake studies because of limited populations (e.g. New Madrid)
or coastal locations (e.g. Charleston and Lisbon). Standard
practice in such cases is to assume isoseismal symmetry, which
in the case of New Madrid would mean doubling the eastern
areas. Doubled areas are listed in Table 3, but also listed are
total areas computed using western reduction factors (see
below), derived from the isoseismal patterns of more recent
New Madrid earthquakes with western coverage.
The two largest New Madrid zone earthquakes since 1812
were 1843 Marked Tree, Arkansas, near the southern end of
the microseismicity zone of Fig. l(a), and 1895 Charleston,
Missouri, near the northern end. [No instrumental records are
known for the 1895 event (Street, Couch & Konkler 1986; Abe
1994; Abe, personal communication 1995).] Regression analysis similar to that described below for New Madrid yields M =
6.3 0.29 for Marked Tree and M = 6.6 0.29 for Charleston,
Missouri (Table 4b). These estimates derive in part from the
isoseismal maps shown in Fig. 6 and in part from other maps
(Nuttli 1974; Street, Couch & Konkler 1986 Stover & Coffman
1993). The east-west isoseismal asymmetry apparent in Fig. 6
is apparent in these other maps as well. The significantly
reduced areas to the west relative to the east would have been
difficult to predict directly from the crustal Q contours of
Fig. 4. Nevertheless, the total measured Ai in all maps are
smaller by factors of 0.62-0.86 than the A i obtained by
doubling the area east of the -north-south lines added to
Fig. 6. These are the western reduction factors used to obtain
the reduced New Madrid Ai of Table 3.
The ENA regressions (eqs 4) are strictly applicable only to
North America east of 90"W latitude (the Mississippi River
provides a convenient approximate boundary) because all but
three of the SCR North America Mo-Ai data points [94 per
cent of the Part 11, Table A3(a) data set] are located in eastern
(ENA) crust. The ENA data are dominated by events from
the high-Q north-central US and eastern Canada, which results
in large A irelative to average SCR crust. Therefore, the New
Madrid eastern A iwith full doubling are appropriate for use
with the eqs (4) ENA regressions but are expected to yield
overestimated M , with the eqs ( 1 )SCR regressions. Conversely,
if the western reduction factors are applied, the SCR regressions
are more appropriate and the ENA regressions should underestimate M,. Therefore, application of eqs (1) with westernreduced A i and eqs (4) with fully doubled Ai should yield
approximately the same results and will serve as cross-checks
on each other.
The 1811 December 16 (02: 15 local) principal event
(NM1)
Based on the local reports of riverboat travellers and settlers,
NM1 probably ruptured a fault outlined by the Bootheel
lineament (Johnston & Schweig 1996) or the southwestern
0 1996 RAS, GJI 126, 314-344
Seismic moment assessment of earthquakes-Ill
323
Figure4. The Stover & Coffman (1993) adaptation of the Nuttli (1981) New Madrid isoseismal map for the 1811 December 16 NMI event
superposed on the coda Q( 1 Hz) contour map of Singh & Herrmann ( 1983). The eastern New Madrid A,, isoseismal contour is extrapolated to
the west (dashed) for comparison with the Stover & Coffman (1993) felt limit and MMI IV isoseismal areas of the M 7.7-7.9 1906 San Francisco
earthquake. The original Singh & Herrmann contours have been extrapolated in places where contour limits are dashed. The L-symbol in eastern
Tennessee marks the location used in Fig. l l ( b ) of reported (Nelson 1924) but unverified rockfalls.
trend of current seismicity (Nuttli 1983) (Figs l a and 12a). Its
source zone is therefore close to the 1843 Marked Tree event
but lies 100-200km to the south of the 1895 Charleston,
Missouri, event. Therefore I used the 1843 event to obtain
western reduction factors for both NM1 and its large aftershock. The main limitations of the NM1 intensity data (Fig. 5a)
are a lack of reports to the west and control on the felt limit
elsewhere. There are undocumented or poorly documented felt
reports from points significantly beyond those reported by
Nuttli (1973, 1981), Street (1980, 1982, 1984), Street & Nuttli
(1984) and Stover & Coffman (1993), for example in
Yellowstone ( Wyoming/Montana), Cuba, Mexico, and Quebec
City, Canada. Documents from parts of Canada, Spanish
Florida, Texas and Mexico have never been systematically
searched, so a constrained felt area may yet be possible with
further historical research.
The reduced Ai of Table3 were evaluated using the SCR
F94 regressions (eqs 1)through A,,,,; the ENA F94 regressions
(eqs 4) were applied to the fully doubled eastern Ai through
A,,. Results are summarized in Table 4(a). As anticipated, the
poorly determined Awl, yielded predicted log(M,) values significantly lower than all other Ai, even when treated as A,,,.
Prediction uncertainties for both log(M,) and M follow Part
I1 for SCR and Table 2( b) for ENA. Best-weighted averages
0 1996 RAS, GJI 126, 314-344
and final uncertainties were computed following eqs (2)-( 3).
Both the SCR and ENA analyses with Afel, omitted yield an
NM1 Mbest of 8.1, but the ENA prediction uncertainty is
about double the SCR value.
A critique of the judgments made in the NM1 regression
analysis begins with the question of its felt area. If the Stover
& Coffman (1993) felt limit were included as Afelt, the M,,,
8.1 would reduce to 7.7 or, as A,,,, to 7.9. I consider the
Stover & Coffman (1993) Afel, to be limited by data reporting
rather than to represent a true felt area; therefore the proper
treatment is to omit it. Next, I selected the F94 functional
regression form over others such as the quadratic developed
in Part I1 and Johnston (1994), because it is based on physical
principles of seismic wave propagation. A final judgment was
that the ENA F94 regressions on fully doubled eastern areas
should not supersede the SCR F94 regression results using the
reduced A i because of the lack of ENA M,-Ai data at large
Ai.They serve as a useful internal consistency check, however,
and it is encouraging that the NM1 Mbestvalues of Table4
are so similar. The best estimate for the size and uncertainty
of NM1, then, is from the SCR F94 regressions using the
Table 3 adopted AIrAVIII values reduced by calibration with
the 1843 event: Mfinal=8.1k0.31 [M,= 1.8 (0.60-5.3) x loz8
dyne cm] (Table 4).
-
-
324
A. C. Johnston
THE 1811-1812 NEW MADRID SEQUENCE
,
0
,km,
300
Figure 5. Intensity data for the 1811-1812 New Madrid sequence from Street & Nuttli (1984). Data points are modified Mercalli (MMI) scale; a
small+after the number indicates an intensity level between it and the next higher MMI level. (a) Intensity data for NM1, 1811 December 16
(02: 15 local). Superposed isoseismal contours are from Nuttli (1981), which are based on a slightly different intensity data set. (b) Intensity data
for NM2, 1812 January 23. Large preceding+or - indicate, respectively, an intensity above or below that of NM1 at the same location. (c) Intensity
data for NM3, 1812 February 7. Symbols are the same as in (b). (d) Isoseismal data for the major aftershock, 1811 December 16 (08: 15 local).
Symbols are the same as in (b).
0 1996 RAS, GJI 126, 314-344
Seismic moment assessment of earthquakes-Ill
Table 4. Regression results for New Madrid.
(a) NM1: 16December 1811 (0215 local)
Log&) +doubledAi *reduced Ai
RegressionType
MMI
-
felt (1-11)
-
-
6.62
6.78
-
-
6.47
felt (111)
IV
V
VI
VII
VIII
6.62
6.78
6.64
-
-
Log(M, ) , 'a
;
f 0.594
M
0;
6.79
7.15
f0.396
f 0.337
7.42
7.80
f 0.281
f 0.401
SCR F94
ENA F94
26.23
26.78
SCR F94
ENA F94
27.18
27.75
f 0.602
SCR F94
ENA F94
27.72
27.92
f 0.764
7.78
7.91
f0.375
f 0.510
f0.506
f0.421
f0.562
6.31
6.48
-
SCR F94
ENA F94
28.52
28.91
f 0.831
f1.484
8.31
8.57
f0.554
f0.989
-
8.30
8.66
f0.501
f 1.230
6.10
6.21
-
SCR F94
ENA F94
28.50
29.05
k 0.752
f 1.845
SCR F94
ENA F94
28.79
~-
f 0.896
-
8.50
f 0.597
SCR F94
ENA F94
28.43
f 0.739
8.25
f0.493
-
5.72
5.83
-
-
5.12
5.23
-
-
--
-
-
-
OPTIONS FOR BEST WEIGHTED VALUES
(i) All Aj with Afelt, without A I ~
(not used-for comparisononly)
SCR F94
ENA F94
27.80
27.34
f 0.29
f 0.40
7.83 f0.19
7.53 f 0.26
(ii) All Ai with AIII, without Afelt
(not used-for comparison only)
SCR F94
ENA F94:
27.86
27.98
f 0.26
f 0.44
7.87 f0.17
7.95 f0.29
SCRF94:
ENAF94:
28.27
28.23
40.33
kO.64
8.14 f0.22
8.12 f0.43
( i v ) LiquefactionSeverity Index (upperbound):
Maximum landslide distance (upperbound):
29.39
> 28.05
-
8.9
> 8.0
-
-
Preferred value [from (iii) SCR & ENA F941:
28.25
f 0.33
8.1
f 0.22
28.25
k 0.14
f 0.47
(iii) AIV - AVIIIonly
AN - A w only
Systematic error contribution (eq. 18, Part U):
NMk Final Value and Uncertainty:
-
f0.09
M8.1 f 0.31
eastern Ai of Figure 4 doubled to the west
*western portion of Aj reduced by correction factor from 1843event (Figure 6b)
(b) New Madrid Summary. 1811 - 1895
sources
Event
N M l l 6 December 1811
(-0215 local)
Used
N73, N81
SC, SN
Approx. No.
MMIdatapts. LogUM,)
Mfid
7
(;
IV, V
VI, VII
VIII
70
28.25
f0.47
8.1 f0.31
111, IV
v, VI
VII
40
26.80
f0.46
7.2 f0.31
580, S82
111, IV
50
27.80
f 0.50
7.8
SN
v, VI
IV, V
VI, VII
45
28.00
f 0.49
8.0 f 0.33
felt, IV
70
25.55
f 0.43
6.3 f0.29
16 December 1811
S80, sS2
aftershock (-0815 local) S N
N M 2 23 January1812
(-0900 local)
AiUsed
f 0.33
VII
NM3: 07 February 1812
(-0345 local)
SSO, sS2
05 January 1843
(05- VT)
N74, H83
SN
v, VI
VII, VIII
felt, IV,
> 300
26.00 f0.44
6.6 f0.29
V, VI
VII, VIII
source: HA, Hopper & Algermissen 1980; H83, Hopper, Algermissen & Dobrovolny 1983;
N73, Nuttli 1973; N74, Nuttli 1974; N81, Nuttli 1981; SC,Stover & Coffman 1993;
580, Street 1980; 582, Street 1982; 586, Street, Couch & Konkler 1986;
SN, Street & Nuttli 1984
31 October 1895
(1l:os VT)
0 1996 RAS, GJI 126, 314-344
N74, H A
sS6,SC
325
Seismic moment assessment of earthquakes-Ill
The 1812 January 23 (NM2) and February 07 (NM3)
principal events
The 1811-1812 New Madrid earthquakes were a complex
sequence of very large events that occurred within the space
of eight weeks. The number of events and the compressed time
frame present some analysis problems; for example, many of
the existing accounts give simply a composite description
or are unclear concerning which event is described. NM1,
NM2 and NM3 are widely accepted as the ‘mainshocks’
of the sequence, although Street (1982) includes the
1811 December 16 (08:15 local) shock as one of the major
events, and in one of his last works Nuttli (1987) includes as
a fifth major event an 1811 December 16 (-12:OO local)
shock. The intensity distributions of NM1, NM2 and NM3
and the largest aftershock are shown in Fig. 5 (adapted from
Street & Nuttli 1984). Isoseismal contours of these data have
been published only for NM1, although, on the basis of the
attenuation of intensity to the northeast, Nuttli (1973) and
Street (1982) rated NM3 slightly larger than NM1 and NM2
slightly smaller. New Madrid residents also ranked NM3 the
most severe, the ‘hard shock’, of the sequence.
In Fig. 5(b,c,d), I have added preceding pluses or minuses
on intensity data points that, respectively, are greater than or
less than same-location data points of NM1. Data points of
the same intensity have no preceding symbol. NM2 has twice
as many lower points as higher and NM3 has four times as
many higher points as lower. This would seem to indicate that
the previous ranking of the principal events of the sequence is
at least qualitatively correct. For NM3, the higher points
strongly concentrate in two areas: the Ohio River Valley system
and the South Carolina coast.
I estimated poorly constrained eastern isoseismal areas for
NM2, NM3, and the 1811 December 16 aftershock from the
Fig. 5 data (Table 3). The analysis was carried out in the same
manner as for NMl, except that the 1895 Charleston, Missouri,
earthquake (Fig. 6a) was used to obtain the western area
reduction factors for both NM2 and NM3. This was done
because these earthquakes probably were on the northern arm
and central segment, respectively, of the Fig. l(a) microseismicity pattern (see Fig. 12a), closer to the 1895 than the 1843
epicentre. Placement of NM3 on the central segment is particularly well constrained because of several eyewitness accounts
on that date that are consistent with faulted disruption of the
Mississippi River bed upstream and in the vicinity of the town
of New Madrid (Penick 1981; Johnston & Schweig 1996).
Final estimated log(M,) and M values and uncertainties are
listed in Table 4( b) without the detail of the individual Ai
regressions. As for NM1, ENA F94 regressions were applied
to the fully doubled Ai,and SCR F94 regressions to the
reduced Ai.In terms of moment magnitude the final estimates
= 7.8 0.33; NM3, Mfinal
= 8.0+ 0.33; and the
are: NM2, Mfinal
= 7.2 & 0.31. These values thereDecember 16 aftershock, Mfina,
fore would rate NM3 slightly smaller than NM1, in disagreement with the previous assessments. I propose that the
difference in relative ranking for these two earthquakes may
be explained by the location and unique faulting mechanism
of NM3 rather than by a larger M,.
The model for NM3 most consistent with all accounts and
with the current seismotectonics of the New Madrid zone is a
large thrust earthquake whose rupture zone broke to the
surface immediately west and south of the town of New Madrid
+
0 1996 RAS, GJI 126, 314-344
327
(Johnston & Schweig 1996). It was significantly closer to the
town than the source fault for NM1, which accounts for the
high-intensity local reports. NM1 and NM2 were probably
strike-slip events (Fig. 12a), meaning that, of the principal
events, NM3 had a unique radiation pattern of Love or L,
waves. Bilateral, up-dip rupture on the NM3 fault plane of
Johnston & Schweig (1996) (see Fig. 12a) would direct a
maximum-amplitude Love-wave lobe northeastwards along
the Ohio River Valley and another southeastwards towards
South Carolina, accounting for the high relative ranking of
NM3 at these azimuths. This model is conjectural and only
intended to demonstrate that plausible alternatives exist to the
‘standard’ model of the New Madrid sequence in which NM3
is the largest event.
Source scaling implications
Specification of an earthquake’s size ( M , ) does not uniquely
define its other source parameters. To accomplish this requires
adoption of specific source scaling relations as well as any
constraints imposed by seismotectonic setting. In this section
I do this for the New Madrid sequence, using the Table 4( b)
M , values. With M , specified, the range of permissible source
dimensions, fault slip and stress and strain drops can be
examined within the limits imposed by the crustal structure
and rheology of the Reelfoot rift (Ervin & McGinnis 1975),
the host crustal structure of the New Madrid earthquakes.
Seismic moment constrains source fault area A and average
slip d via the well-known defining equation of scalar moment
M , = p d A (Aki 1972). I use an upper crustal rigidity p derived
from the shear-wave velocity profile in Chiu, Johnston & Yang
(1992). This yields an average rigidity for the upper crust
(depth h= 20 km) of p = 3.5 x 10“ dyne cm-2, similar to that
used by Gomberg (1993), but slightly higher than that used
by Nuttli (1983). Because I shall consider faulting to greater
depths than did those authors, I also derive a lower-crust
rigidity (h=20-33 km) of p=4.5 x 10l1 dyne cm-2. With p
specified, a given M , restricts the product dA but not d or A
separately. This is where constraints from crustal rheology and
source scaling become important.
Perhaps the most problematic source-scaling question is
depth of faulting h, which controls the down-dip width W of
coseismic rupture. Normally, h is taken as the depth of the
transition from brittle to plastic rheology. In quartzofeldspathic (continental) crust, this is controlled primarily by
temperature, and for the quartz component occurs at
300-350°C or roughly h=15-25 km. If the quartz brittleplastic transition (bpt) limits W, then very long faults are
required for M 2 8 earthquakes.
Based on work by Strehlau (1986) and Sibson (1980, 1986),
Scholz (1988, 1990) developed a synoptic shear-zone model of
continental crust that incorporates a temperature-dependent
transition zone between strictly brittle and strictly plastic
rheology. Depending on the temperature-depth profile and
rock composition, the thickness of this intermediate zone
of semibrittle behaviour may be 10 km or more. The upper
boundary bpt is the depth (temperature)at which quartz begins
to deform plastically ( T - 300 “C).The lower boundary, determined by the onset of feldspar plasticity, occurs at T-450 “C.
Brittle rupture to transition-region depths is supported by
both petrologic evidence (Sibson 1986) and analytical studies
(Das 1982).
328
A. C. Johnston
In the synoptic shear-zone model, the bimodal brittle-plastic
crust is replaced by one with three distinct rheological zones:
(1) the brittle zone, beginning at shallow depth, in which
seismic rupture can both nucleate and propagate; (2) the
semibrittle (or semiplastic) zone at depths between the onsets
of quartz and feldspar plasticity, in which rupture cannot
nucleate but can propagate from a nucleation event in (1);and
(3) the plastic zone in which rupture can neither nucleate nor
propagate. Crustal depths corresponding to the transitions
between zones (h, for brittle-semibrittle, h, for semiplasticplastic) are controlled by the temperature gradient of the
region. Fig. 7 shows a temperature-depth profile for the New
Madrid seismic zone computed for a high crustal heat flow of
60 mW m-,, compared with a profile for 'normal' central US
crust with a heat flow of 45 mW rn-, (Liu 1993). I take these
as high and low estimates, respectively, and use an intermediate
profile (dashed line) to estimate h, N 20 km (extremes of
16-27 km) and h, N 33 km (extremes of 25-46 km) for the
New Madrid zone.
The aspect ratio of New Madrid faulting is unknown and
probably not constant, so specification of W does not constrain
the fault length L. This must be done by source scaling;
moreover, with M , and W constrained, there may still be a
trade-off of L with 2. I consider a range of possibilities for
these parameters as expressed in terms of the static strain drop
A&= cd/A or equivalently the static stress drop Aa=cpJ/A,
where c is a shape factor close to 1 and A is a characteristic
rupture dimension (Kanamori & Anderson 1975; Scholz 1990).
The A may be L or W, as in the faulting models of Scholz
(1982, 1990), or r, the radius of an equivalent circular fault.
For the latter case, c=77c/16 and the static stress drop is just
a Brune (1970) stress drop. I consider all these cases (Table 5),
but follow Kanamori (1994) and, for c- 1, take A = r W , the
dimension of an equivalent square fault, as the appropriate
scale length for estimating static stress and strain drops.
-
-
s
20
Faulting scenarios for NM1, NM2 and NM3 are explored
in Table 5. None is unique, but they are intended to illustrate
the range of options in L, and A a that would be available
for given M,, p and W. For N M l , which is taken to be a
vertical strike-slip fault, two cases are considered: ( 1) rupture
width limited by the bpt at h, =20 km; and (2) rupture propagation to the fully plastic limit, hL=33 km. In both cases, and
for other earthquakes studied, d is taken as being evenly
distributed over the available rupture area, although in reality
maximum slip probably would concentrate near but above the
quartz bpt in the zone of highest strain energy density and
taper towards the surface and h2. Similar temperatures and
depths apply to NM2 and NM3, except that NM3, modelled
as a dip-slip fault with dip angle of -32" (Chiu, Johnston &
Yang 1992), uses the high geotherm model of Liu (1993),
which from Fig. 7 limits depth of rupture to h2-25 km and
W to -45 km. This choice was arbitrary; for the moderate
geotherm with h2 = 33 km, 50 km or more down-dip fault width
is possible.
In Table 5 fault models considered more plausible than
others are highlighted in bold type; among these, the boxed
models are illustrated in Fig. 12. The principal criterion for
'plausibility' was that the model should have static stress and
strain drops close to global or intraplate averages: -30-100
respectively (e.g. Kanamori & Anderson
bar and -2 x
1975; Kanamori 1994). A secondary criterion was the difficulty
of envisioning very long faults in the tectonic setting of the
New Madrid seismic zone (Fig. la): segment lengths defined
by current seismicity do not exceed 150 km, although multiplefault-segment involvement in one or more of the principal
event ruptures is possible, even probable.
All except one of the preferred fault models of Table 5 for
the three principal New Madrid earthquakes require faulting
to depths of 25-33 km, that is, rupture into the brittlesemibrittle transition zone. NM1 is modelled as a vertical
strike-slip event with L=120-180 km, W=33 km, 2-8-11 m,
and static stress drop between 40 and 70 bar for a strain drop
of 1.0-1.8 x
The model for the smaller NM2 is vertical
strike-slip, with L =50-100 km, W=20-33 km, 2-5-10 m,
Finally, NM3 is
An= 35-95 bar, and A&=0.85-2.4 x
modelled as a thrust fault, dip -32", L=60-100 km, W =
45 km, 2=6-10 m, Aa=35-70 bar, and A&=0.9-1.9 x
It
is tempting, and also supported by some eyewitness accounts,
to place NM1, NM2, and NM3 on the main microseismicity
trends of Fig. l(a). If one does this (Fig. 12a), the boxed fault
models of Table 5 are favoured.
a
40
THE 1886 CHARLESTON EARTHQUAKE
\
60
0
1 1 1 1 1 1 1 1 1 1 1 1
200
1 1 1 1 1
\,
I I I 1 I
800
Temperature ("C)
400
600
1000
Figure 7. Temperaturedepth profiles for the crust of the New Madrid
seismic zone and the central US (from Liu 1993), using heat-flow
values of 60 and 45 mW m-z respectively. If these are taken as the
high and low thermal gradient bounds, then the dashed profile
represents an average profile for New Madrid. Where this profile
reaches 300°C and 450°C marks h,, the transition from brittle to
semibrittle rheology, and h,, the transition from semiplastic to fully
plastic rheology.
Intensity effects are much better documented for the Charleston
earthquake than for New Madrid. By 1886, the entire central
and eastern United States had been settled and incorporated
into the Union, in contrast to the frontier setting of New
Madrid in 1811. Furthermore, there was a time lag of only
three years for Dutton's (1889)Charleston analysis, in contrast
to the century-long lapse before Fuller (1912) analysed New
Madrid.
Isoseismal regression analysis
Fig. 8, the Bollinger (1977) generalized isoseismal map of the
Charleston earthquake, serves as the base map for the M,-A,
0 1996 RAS, G J I 126, 314-344
Seismic moment assessment of earthquakes-Ill
329
Table 5. New Madrid source parameters.
WOU~) Yequiv(W +F(km)
d(m)
L-model
W-model Brune (Ypn&)
=(bar)
=(bar)
=(bar)
-*-model
G(bar)
NMl, 16 December 1811: log(Mo)= 28.25, M = 8.1
.
.
12, t -W
h = h 2 0 k m b t u k e s h p . vertical fault. u = 3.7 x lQ
300
20
44
77
8.0
10
150
95
40 1.Ox10”1
250
20
40
71
9.6
15
180
125
50 1 . 4 ~ 1 0 ~
200
20
36
63
12.0
20
220
170
70 1 . 9 l@
~
150
20
31
55
16
40
295
265
110 2.9x10”1
100
20
25
45
24
90
445
485
200 5.4x10-4
50
20
18
32
48
355
890
1370
560 1 . 5 10-3
~
. . a crust. h - h7 - 33 kmk vertical fault. u- = 3,9 x lQ11
(ii)
e &s
I
300
250
200
33
33
33
56
51
46
100
91
81
4.6
5.5
6.9
180
160
140
120
33
33
33
33
43
41
38
36
77
73
68
63
7.7
8.6
9.9
11.5
100
80
65
50
33
33
33
33
32
29
26
23
57
51
14
17
21
28
46
41
5
10
15
15
20
25
35
55
85
130
215
55
65
80
90
100
115
135
165
205
250
325
45
60
80
95
115
140
175
230
320
435
645
20
25
35
4.fjX10-5
6.1 x
8 s X10-5
40
45
55
70
1 . 0 ~ 1 0 ~
1.2x10-4
1.5x1@
1 . 8 ~ 1 0 ~
95
130
180
265
2.4~10-4
3.4xlQ4
4.6~10~
6.8x10”1
40
1.0~10~
I
NM2, 23 January 1812: log(Mo) = 27.80, M = 7.8
Expanded w
e (h = hi&-):
strike slip, vertical fault
1
150
20
31
55
5.7
100
20
25
45
8.5
80
60
40
120
20
20
20
33
23
20
16
36
40
35
28
63
10.7
100
65
60
50
33
33
33
33
33
32
29
26
25
23
57
51
46
44
41
4.9
6.1
7.5
8.2
9.8
40
33
20
36
12
80
14
21
4.1
15
30
50
90
195
15
20
30
45
55
75
120
105
160
195
265
395
50
60
70
90
95
115
145
95
170
240
370
680
60
80
115
155
175
230
320
70
1.9~10~
100
150
280
25
2.7~10~
4.1~10~
7.5x10”1
6 . 5 10-5
~
35
45
65
70
95
13.5~10-5
1.2x10-4
1 . 6 x d
1 . 8 ~ 1 0 ~
2 . 4 ~ 1 0 ~
130
6.0~10~
I
NM3,07 February 1812: log(Mo) = 28.00, M = 8.0
W a n d e d Rupture. hieh geotherm (h1=16 km. h2 = 25 km): dip-slip, dip -32”, ,
ii
= 3.7 x loll dyn/cm2
160
45
48
85
3.8
10
30
40
15 4.4x10-5
140
45
45
79
4.3
10
35
50
20 5 . 4 ~ 1 0 - 5
120
45
41
73
5.0
15
40
60
25 6.8 x
100
45
38
67
6.0
20
50
so
35 9. 0~ 10- 5
90
45
36
64
6.7
25
55
95
40 1.ox 10-4
40
65
125
50 1 . 4 ~ 1 0
45
33
58
8.0
[ 75
60
45
29
60
52
10.0
80
175
70 1 . 9 ~ 1 0
50
45
27
90
47
12
100
230
95 2 . 5 ~ 1 0
45
45
25
110
45
13
110
265
110 3 . 0 ~ 1 0
32
45
21
38
19
215
155
445
185 4 . 9 ~ 1 0
-bold
I~
~
~
~
~
indicates fault parameters considered more physically reasonable than the others; boxed modelssee Fig. 12
- stressvops rounded to nearest 5 bars
regression analysis. It is a modified Mercalli reinterpretation
of the data for Dutton’s original Rossi-Fore1 map. The isoseisma1 contours are based on > 1000 intensity observations (plus
an additional 235 ‘not felt’ reports), each one a consensus
value of three independent analysts (Bollinger & Stover 1976).
Except for a small fraction of the Afel, isoseismal, the entire
map is in the eastern United States; therefore, the best
regression application should be ENA regressions using fully
doubled onshore areas. As was the case for New Madrid, SCR
regressions on these doubled areas would be expected to
overestimate M,, but, unlike for New Madrid, there are no
other events that can be used to calibrate an Ai reduction.
A log(M,)-log@,) regression analysis of Charleston is not
0 1996 RAS, G J I 126, 314-344
straightforward because of complicating factors concerning the
high-intensity isoseismals and the felt area. Chapman et al.
(1990) have argued, as did Dutton (1889), that the tapering
edge of the wedge of Atlantic Coastal Plain sediments amplified
Charleston’s strong ground motion at inland sites near the
Fall Line, the Coastal Plain-Piedmont province contact. This
effect could have produced abnormally large isoseismal areas,
affecting primarily A,,, and A,,,,. I therefore consider two
alternatives for Charleston, the first a standard regression
analysis, the second taking the sedimentary wedge effect into
account for A,,, and Av,,, (but not for A,,, whose perimeter is
well inland of the Fall Line).
The Charleston Ais from Fig. 8 are listed in Table 3. As with
330
A. C. Johnston
950
900
850
800
750
700
65O
500
450
35'
30
Figure 8. The generalized Bollinger (1977) isoseismal map (modified Mercalli) for the 1886 Charleston earthquake. The map is based on data
points (not shown) reinterpreted from Dutton (1889), who used the Rossi-Forel scale, and are consensus values assigned by three different analysts.
The L-symbol added in western North Carolina marks the landslide location (Dutton 1889) used in Fig. 11(b). The MMI 11-111 contour should
be 1-111 ( G .Bollinger, personal communication, 1995).
New Madrid, nearly 50 per cent of areas had to be inferred
by symmetry, in this case because of coastal location rather
than lack of population. The Bollinger (1993) values in Table 3
for AVrAv,,, are onshore values, purposely not doubled in
order to compensate for the sedimentary wedge effect.
Charleston's felt limit should be much better determined than
that of New Madrid because it falls within the well-settled
American midwest and because Charleston occurred in the
0 1996 RAS, GJI 126, 314-344
Seismic moment assessment of earthquakes-Ill
mid-evening hours (21 :50 local) as opposed to 0215 local for
NM1, yet it exhibits the same problem as NM1: treatment of
AI,,, as Afelt yields a predicted M , anomalously low relative
to M,, from all other A i [case ( i ) , Table 61. However, treated
as A,,, [case ( i i ) ] , it yields results consistent with the other
regressions. Dutton ( 1889)indicates that Charleston was lightly
but clearly felt in New Orleans, Boston, central Missouri,
Cuba, and Bermuda, points from 80 to >300 km beyond the
Fig. 8 felt limit. Dutton himself (p. 333) ‘roughly approximates’
the felt area as 2.5 million square miles (6 560 000 km2), 72 per
cent larger than the Table 3 dfelvBased on the large number
of ‘not felt’ reports, Bollinger (personal communication 1995)
considers that the A,,,, contour of Fig. 8 is appropriately Afelt,
with the more distant points as outliers.
A range of Charleston regression alternatives is included in
Table 6: regression results using both SCR F94 and ENA F94
formulas; AfeIt, A,,, and an average regression using the
Charleston A,,,, of Fig. 8; and A,,, and A,, regressions with
both doubled (this study) and non-doubled (Bollinger 1993)
areas. The best-weighted values of Table 6 take various combinations of these regressions so that comparisons of the effects
of various assumptions are possible. For example, treating the
felt limit as Afelt gives a predicted Mfelt=6.5-6.7, but, when
treated as A,,,, an M of 7.2-7.3 is obtained. Since this outermost
isoseismal carries the lowest uncertainty, it exerts a large
influence on the best-weighted M-value, Mbest.Use of the Afelt
regression yields Mbest= 7.0-7.6; use of A,, yields Mbest=
7.3-7.7; and omission of the felt limit altogether yields Mbest=
Table 6. Regression results for Charleston.
MMl
Log(Ai) Regression Type
(i) felt (1-11)
6.58
SCR F94
ENA F94
26.08 f 0.593
25.84 f0.468
6.68 f0.396
6.53 f0.312
(ii)felt (111)
6.58
SCR F94
ENA F94
27.01 f 0.417
26.78 f0.482
7.31 f0.278
7.15 f0.321
6.58
SCR F94
ENA F94
26.54 f 0.60
26.31 f 0.48
7.00 f0.40
6.84 f0.32
IV
6.45
SCR F94
ENA F94
27.62 f 0.560
26.99 f 0.668
7.71 f0.374
7.30 f0.445
V
6.19
SCR F94
ENA F94
27.94 f 0.782
27.35 f 1.036
7.93 f0.521
7.54 f0.691
VI
5.73
SCR F94
ENA F94
27.43 f 0.661
27.28 f0.925
7.59 f0.441
7.48 f0.617
VII
5.47
SCR F94
ENA F94
SCR F94
28.06 f0.759
8.01 20.506
27.41 f0.689
7.57 f0.459
SCR F94
ENA F94
SCR F94
28.00 f 0.642
_ _ 27.35 f0.554
7.97 f0.428
_ _ 7.54 f0.369
(iii)felt (1-111)
aver. (i & ii)
*5.19
VIII
4.96
*4.66
-
-
- -
OPTIONS FOR BEST WEIGHTED VALUES
( a ) All Ai: with
(i), without (2) and (iii)
SCR F94 (with AVII-AVIII): 27.44 f 0.27
SCR F94 (with *AVII-*AVIII):
27.25 f 0.26
26.48 f0.34
ENAF94:
7.59 f0.18
7.47 50.17
6.95 f0.22
(b) All Ai: with (ii), without (i) and (iii)
SCR F94 (with AVII-AVIII):
27.52 f 0.24
SCR F94 (with *AVII-*AVIII): 27.36 f 0.23
ENAF94:
26.96 f0.34
7.65 f0.16
7.54 f0.16
7.28 f0.23
(c) All Ai: with (iii), without ( i ) and (ii)
SCR F94 (with AVII-AVIII): 27.54 f 0.27
SCR F94 (with *AVII-*AVIII): 27.35 f 0.26
ENAF94:
26.73 f 0.34
7.66 f0.18
7.53 f0.17
7.12 f0.23
SCR F94 (AIV- AVIonly):
27.63
ENA F94 (AIV- AVI only):
27.15
(e) ENA F94 from (a) & SCR (*AvII-*AvIII): 26.82
ENA F94 from (b) & SCR ~Av~I-*AvIII): 27.12
ENA F94 from (c) & SCR (*AvII-*AvIII): 26.98
(d 1
(R
Liquefaction Severity Index:
Max.landslide distance (upper bound):
Preferred Value [from (e)]:
Systematic error contribution (eq. 18, Part 2):
Charleston: Final Value and Uncertainty:
f 0.38
f 0.48
f 0.26
f 0.27
f 0.27
27.15 (f 0.53)
>27.45 27.00 f 0.27
27.00
f 0.11
f 0.38
7.72 f 0.25
7.40 f0.32
7.18 f 0.18
7.38 f0.18
7.29 f0.18
7.4 (f0.35)
>7.6 __
7.3
f0.18
f 0.08
M 7 3 f 0.26
* Bollinger (1993):AWI,AVIIIcorrected for wedge effect of coastal plain sediments
() qualitative uncertainty bounds
0 1996 RAS, G J I 126, 314-344
331
332
A. C . Johnston
7.4-7.7. The Mbestfrom the ENA regressions averages 0.3-0.6
M units lower than from the SCR regressions. The Dutton
Afelt yields M 7.3 with either SCR or ENA regressions.
The largest Charleston Mbestof 7.7 results from application
of the SCR F94 regressions with A,,, omitted and with full
doubling on AvII and A,,,,. The smallest, an Mbestof 7.0, is
obtained with the ENA regressions for Afelt-AVI,and A,,, and
A,,,, omitted. Use of the Bollinger (1993) reduction of A,,,
and AvIlr in eqs (1) reduces Mbestby 0.13 M units. The
final Charleston size estimate is taken from option (e) of
Table6 and makes maximum use of the available intensity
= 7.3 -t 0.26; M , = 1.0
data. The preferred estimate [Mfinal
dyne cm] ( 1) averages treatment of the felt
(0.42-2.4) x
limit as Afel, and A,,,; ( 2 ) applies the eqs (4) ENA regressions
up to A,,; and (3) adopts the Bollinger (1993) AvII and AvIII
for the eqs ( 1 ) SCR analysis. The Mfinal=7.3 is in good
agreement with the M 7.4 obtained from liquefaction data in
a following section.
Charleston source scaling
No direct geological or geophysical evidence for the source
fault of the Charleston earthquake has been uncovered despite
large and concentrated efforts (e.g. Rankin 1977, Gohn 1983,
Dewey 1985), although Marple & Talwani (1993) have identified a -200 km long, NNE-trending zone of subtle topographic highs and drainage anomalies that crosses the
Charleston epicentral region (Fig. lb) and could be the surface
expression of a fault at depth. Current seismicity and relict
liquefaction features (Fig. lb) mostly are contained within the
-40 x 50 km MMI X meisoseismal area. A 200 km fault length
is much greater than a mid-M 7 requires; therefore, the
meisoseismal zone is the best, albeit indirect, constraint on
Charleston's source dimensions other than its seismic moment.
The crust in the Charleston region is 30 km thick (Luetgert,
Benz & Madabhushi 1994), compared to -40 km in the New
Madrid vicinity. There are no computed temperaturedepth
profiles for the Charleston area as there were for New Madrid,
but there are sparse heat-flow data (McCartan & Gettings
1991) that average 53mW rn-, over inferred magmatic
intrusions and 43 mW m-2 elsewhere. Given this heat-flow
variability, a temperature of 300 "C could be reached anywhere
from 15-26 km depth using Fig. 7 as a proxy for Charleston
geotherms. Therefore I model the Charleston source fault to
depths of 16 km (case i ) or 25 km (case ii) (Table 7a). Note that
these are depths corresponding to h,, the brittle-semibrittle
transition depth. For a mid-M 7 earthquake there is no need
to invoke rupture into the semibrittle regime (to h, depth),
although neither would it be prohibited. The focal depths of the
12 km (Shedlock 1987),
current microseismicity extend to
suggesting that case (i) may be more appropriate.
Shear-wave velocities and densities from the Luetgert, Benz
& Madabhushi (1994) coastal plain crustal model yield an
average rigidity of p=3.7 x 10" dyne cm-2 for the Table 7(a)
3.9 x 10" dyne cm-, for case ii. With these
case i and ,!i=
rigidities and with down-dip fault width W approximated by
hl, a range of faulting models for an M 7.3 Charleston
earthquake is generated and listed in Table 7(a), with the more
plausible models in bold type. Fault length L is varied from
20 to 180 km, and stress and strain drops are computed as for
Table 5. A representative fault model with a static stress drop
of -50 bar is selected for each case (Fig. 12b), which yields a
fault length of 50 km (case i) or 30 km (case ii), both with
-
-
-
average slip of 3.4 m. Case ( i i ) may be less likely because of
the abundance of warm intrusions in the region. The maximum
regional tectonic stress for Charleston is compressional and
oriented northeast (Zoback 1992)or 30" to the NNE tectonic
zone of Marple & Talwani (1993). If the Charleston earthquake source fault is subvertical and aligned on this zone, a
right-lateral strike-slip mechanism is indicated (Fig. 12b).
-
THE 1 7 5 5 L I S B O N EARTHQUAKE
The Lisbon earthquake presents difficulties, principally in
sorting out incomplete, often conflicting, information regarding
intensity distribution, origin time, and timing of the strong
shaking and tsunami arrivals. Fig. 1(c) shows the bathymetry
and tectonic elements of the offshore Portuguese continental
margin along with four Lisbon epicentres from the literature.
For isoseismal mapping purposes, the epicentre of Machado
(1966) at 36.45"N, 11.25"W will serve. It places the earthquake
on the south flank of Gorringe Ridge, an uplifted section of
oceanic lithosphere (Purdy 1975), 70 km NW of the epicentre
of the tsunamigenic M 7.8 1969 St Vincent earthquake and
-320 km SSW of Lisbon. Machado's epicentre was based on
a general consideration of intensity distribution. Epicentres
estimated from tsunami arrival times (Milne 1841; Reid 1914)
are closer to Lisbon (Fig. lc). All authors acknowledge large
uncertainties because of poorly documented data, especially in
regard to the timing of the seismic-wave and tsunami arrivals
at various locations.
-
Isoseismal regression analysis
The intensity distribution of the Lisbon earthquake has been
documented in detail only on the Iberian Peninsula (Pereira
de Sousa 1919; Machado 1966; Martinez Solares, L6pez Arroyo
& Mezcua 1979;Johnson & Kissenpfennig 1977). More general
summaries of the far-field intensities can be found in Bevis
(1757), Reid (1914), Davison (1936) and Richter (1958). These
sources differ mainly in the degree of scepticism applied to the
reports. For example, felt reports from North America (e.g.
Woerle 1900), which evidently confused Lisbon with the M > 6
New England earthquake of 1755 November 18, are discounted
effectively by Reid (1914). Reid and Richter (1958) do not
accept felt reports from variom sources for Denmark,
Scandinavia, Iceland, Egypt, and Greenland. Richter discusses
the difficulty of separating ultra-long-period effects of large
earthquakes, such as seiches and pendulum motion, from true
felt effects.
Fig. 9 presents the isoseismal area map used for Lisbon's
regression analysis. It is derived mainly from Machado (1966),
who depended heavily on Reid ( 1914). Its felt area is conservative, rejecting the felt reports at extreme distances but accepting
the felt reports near Hamburg and in the Cape Verde Islands
and intensities of -111 for the Azores, -1V for the Canary
Islands, and -V for Madeira. Fully one-half of this suite of
isoseismal areas is in oceanic crust, and about 20-30 per cent
lies in active tectonic crust (see Fig. lo), a distribution that
calls into question the applicability of the SCR regressions.
Fortunately, there is an isoseismal map (although no Afelt)for
the 1969 St Vincent event (Lopez Arroyo & Udias 1972) of
known M , that can be used to calibrate a Lisbon isoseismal
analysis. Fig. 10 displays a same-scale comparison of the
Lisbon and St Vincent maps and Table 3 lists the measured A,.
The extent of Lisbon's isoseismal pattern in oceanic crust
raises the question of attenuation of seismic waves over oceanic
0 1996 RAS, GJI 126, 314-344
Seismic moment assessment of earthquakes-III
333
Table 7. Charleston and Lisbon source parameters.
(a) Charleston: log(Mo)= 27.00, M = 7.3
L&m)
L-model
ho(bar)
W(km) T,quiz&m) m
(
k
m
)
d(m)
W-model Brune f Y + . , , h )
=(bar)
=(bar)
7
=(bar)
(i) Brittle crust, hl= 16 km: vertical fault, E = 3.7 x 10l1 dyne/ax?
1
160
14G
120
100
16
16
16
16
29
27
51
47
25
23
44
80
60
50
20
17
16
14
12
36
31
28
25
30
16
16
16
16
16
20
16
10
40
1.1
1.2
1.4
1.7
40
2.1
2.8
3.4
4.2
5.6
22
18
8.4
(ii) Brittle crust, hi= 25 km: vertical fault,
1
160
140
120
100
80
25
25
25
25
25
36
33
31
28
25
60
50
40
30
25
25
25
25
25
25
20
25
63
59
45
22
20
18
15
14
39
35
32
27
25
1.7
2.1
2.6
3.4
4.1
13
22
5.1
55
130
195
80
110
150
230
425
2
2
3
4
5
10
15
25
45
65
100
10
10
15
15
20
25
30
40
55
65
80
10
10
15
20
25
40
55
75
120
155
220
100
z=3.9 x 10l1d y n / a n 2
0.6
0.7
0.9
1.0
1.3
50
25
30
35
40
50
65
80
2
3
5
5
10
20
25
40
70
155
20
25
30
40
55
5
10
10
15
2.1 x 10-5
2.6 10-5
3.2~
10-5
4.2~
20
35
45
60
95
5.9 x 10-5
9.1~10-5
1.2XlO~
1 . 7 ~ 1 0 ~
2 . 6 ~ 1 0 ~
I
175
4.7x 10-4
4
5
5
1 . 010-5
~
1.zX10-5
1 . 6 10-5
~
2.1 10-5
2 . 9 10-5
~
10
10
15
25
30
50
65
90
4.4~104
5.8 x lod
8.1~lO-~
I.ZXIO*
J
1.6~lo-*
2.3~10~
(b) Lisbon: log(Mo) = 29.10, M = 8.7
Oceanic lithosphere, 600" isotherm hi
50 km: dip-slip fault, d i p -40",
-
L(km)
W ( h ) Yequidkm)
500
400
350
1
300
250
230
200
180
150
100
80
80
80
80
80
80
80
80
80
80
m(h)d(m)
113
101
94
200
87
80
77
71
68
155
141
136
126
120
110
62
50
179
167
89
L-model W-modd
4.8
6.1
6.9
8.1
9.7
10.5
12.1
13.5
16
24
ji= 6.5 x 1011dyn/anz
T
=(bar)
ba(bar)
ba(bar)
5
10
15
15
25
30
40
50
70
155
40
50
55
65
80
85
100
110
130
195
40
55
65
80
110
125
150
175
235
430
. ~-liLW-moi~
ha(bar1
15
3
2 . 4 lo5
~
3.4x 10-5
4.1 10-5
35
45
50
60
75
5.zX10-5
6.8~
7.8~
9 . 6 l~o 5
1.1x10-4
20
95
175
-bold
indicates fault parameters considered more physically reasonable than the others; boxed model- stress rops rounded to nearest 5 bars
paths at the predominant felt and damage frequencies that
100-150 km, L,
control intensities. In SCR crust, beyond
surface waves have the largest amplitudes; their geometrical
spreading and attenuation were used by Frankel (1994) in
developing the F94 regression functional form. In oceanic crust
S, (or So), trapped shear waves propagating entirely in oceanic
lithosphere, are the analogue to continental L, and in fact
convert efficiently to L, at continental margins. Across Europe
and North Africa, continental Lg was probably the principal
strong-ground-motion phase, at least for MMI VII, even
though the Lisbon hypocentre was in oceanic lithosphere. The
oceanic S , phase should have been the principal phase sensed
on the oceanic islands Madeira, Canary, Azores and Cape Verde.
Oceanic lithosphere is a very efficient high-frequency
(> 10 Hz) S , wave guide but is only moderately so at 2-4 Hz,
the frequency band of greatest human sensibility (Frankel
1994). For example, Sereno & Orcutt (1987) find that shear-
-
Q 1996 RAS, GJI 126, 314-344
I
1 . 5 lo4
~
lo4
2.7~
Fig. 12
wave Q decreases from >2000 above 10 Hz to 300-500 at
2-3 Hz, and Butler et al. (1985) measure an S,, Q of 500 at
2.5 Hz and 3000 at 22 Hz. The 2-3 Hz anelastic attenuation
of oceanic lithosphere is comparable to L, attenuation in SCR
crust, for example the 1 Hz L, coda Q of 350-450 obtained by
Canas et al. ( 1987)for the SCR portion of the Iberian peninsula.
Therefore, for frequencies that control intensity levels, the
attenuations of S, and L, can be considered the same, even
though the seismic phases, and therefore geometrical spreading,
are different. However, the presence of substantial active
tectonic continental crust with L, coda Q<200 (Canas et al.
1987)and the relatively high attenuation of oceanic lithosphere,
Iberia, France and much of southern Europe in comparison
to average SCR crust leads to the expectation that Lisbon's
predicted M , from SCR Ai regressions will be low. Thus
calibration with the 1969 St Vincent earthquake isoseismal
pattern is important.
334
A. C . Johnston
4 Oo
30a
20'
,J
loa
f
i--J
I
1000 km
I
Figure 9. Generalized isoseismal map developed in this study for the 1755 Lisbon earthquake. Isoseismal contours are taken primarily from
Machado (1966) with control on the Iberian peninsula from Johnson & Kissenpfennig (1977) and Martinez Solares, Lopez Arroyo & Mezcua
( 1979). The felt limit discounts poorly documented felt reports from Iceland, Fennoscandia, and Egypt. Important isoseismal control points are
labelled A, Azores; C, Canary Islands; CV, Cape Verde Islands; H, Hamburg; L, Lisbon; M, Madeira Islands.
Table 8 contains the regression results for both the St
Vincent and the Lisbon earthquakes. The St Vincent best= 7.35 and
weighted averages with and without A,,, are Mbest
7.66, respectively, compared to M 7.82 (Fukao 1973) from
long-period surface-wave equalization. Since the percentage of
active and SCR crust differed for Lisbon's Afel, and A,,, versus
St Vincent's A,,,, I averaged the two Mbestvalues to obtain M
7.5 for St Vincent. Thus isoseismal data predict a moment
magnitude for St Vincent 0.32 M units smaller than its
instrumental value. This is the calibration factor for Lisbon,
and it is consistent with the low-Q crustal environment of the
St Vincent isoseismal areas.
The Lisbon isoseismal map of Fig. 9 yielded the Ai measurements listed in Table 3. As mentioned, the felt area is conservative, yet at > 14 million km2 (effective radius -2200 km or
-20°A), it is the largest documented felt area of all the world's
shallow earthquakes because the vast majority of M > 8 events
occur in low-Q plate boundary and active tectonic crustal
environments. [Some great deep-focus events, notably the 1994
June 9 M 8.3 Bolivian earthquake, have been felt to A>60"
0 1996 RAS, GJI 126, 314-344
Seismic moment assessment of earthquakes-Ill
(b) 1969 St. Vincent
I
8O
335
/
I
4O
Figure 10. Isoseismal comparison of the (a) 1755 Lisbon and (b) 1969 St Vincent ( M 7.8) earthquakes. (a) is from this study and (b) from Lopez
Arroyo & Udias (1972). The Betic, Atlas, and Pyrenees mountain belts (shaded) are active tectonic, not SCR, crust. The set of St Vincent isoseismal
areas can serve as calibration for the determination of the Lisbon earthquake's seismic moment from Fig. 9 isoseismals assuming the two events
are nearly co-located and have similar mechanisms and stress drops. The L-symbols in (a) denote the landslide locations from Martinez Solares,
Lopez Arroyo & Mezcua ( 1979) that are used in Fig. 11 (b).
(e.g. Anderson, Savage & Quaas 1995).] The ranking for
Lisbon must be tempered by the fact that much of its felt area
is inferred or extrapolated because of its oceanic location.
Results of an eqs (1) SCR regression analysis on individual
Lisbon Ai range from a low of M 8.1 for Avl to a high of M
8.7 for Afel,. The Lisbon F94 Mbestis 8.4; with the St Vincent
calibration, Mhst increases to 8.7. Having the full complement
of seven A ifrom Awl, to Avl,, yields a small weighted-average
uncertainty of 0.17 M units. However, when calibration-factor
uncertainty and systematic error are included, the final Lisbon
M+oG are: Mfi,,,=8.7f0.39
[M,= 1.26 (0.32-4.9) x loz9
dyne cm].
Lisbon source scaling
The Lisbon earthquake nucleated in oceanic lithosphere, which
is olivine-rich and maintains brittle behaviour to considerably
higher temperature than continental crust. The thermal regime
of oceanic lithosphere correlates tightly with its age. In a model
of oceanic lithosphere as a slab cooled by vertical conduction
only, isotherm depth h is a function of lithospheric age t = x / v :
baa, where x is distance from the spreading axis and o is
the half-spreading rate. In depth profile, the isotherm is parabolic, deepening with distance (age) from the ridge axis. Several
0 1996 RAS, G J I 126, 314-344
studies of oceanic intraplate earthquakes have found that maximum hypocentral depth increases with t and seems to be limited
by isotherms in the 600-800°C range (e.g. Wiens & Stein 1983,
Bergman 1986) in contrast to 300-450 "C for continental crust.
Oceanic lithosphere off the Portuguese continental margin
is old, 130-170 Myr. Grimison & Chen (1986) modelled the
lithosphere of this region and found that temperatures of
600-800°C occur at depths of 40-70 km. In addition, from
depth phases they determined a focal depth of 50+5 km for a
1969 May 5 St Vincent aftershock. Therefore, a lithospheric
depth of at least 50 km and a down-dip fault width W of at
least 80 km should be in the brittle regime for a dip-slip
( d i p s 40") source fault. An average rigidity for the upper 50 km
of oceanic lithosphere is F = 6.5 x 10" dyne cm-'. Fukao (1973)
used a hypocentral depth of 33 km and p=6.0x lo1'
dyne cm-' for modelling the St Vincent earthquake.
Table 7(b) gives a range of possible faulting models of the
Lisbon earthquake in the format of Table5, using the constraints in M,,, W, and ji described above. As fault length L
varies from 500 to 100 km for constant W=80 km, average
slip increases from 4.8 to 24 m and static stress drop increases
from 15 to 175 bar. The greater rigidity of oceanic lithosphere
compared to continental crust means that smaller fault dimensions and average slip are needed in the oceanic setting to
-
336
A. C . Johnston
Table 8. Regression results for Lisbon.
Log(Ai)
MMI
Regression Type
Log(&)
ego
M
a&
1969 St. Vincent
cart-e:
-
-
111
6.42
SCR F94
26.39 f0.411
6.90 f 0.274
IV
6.26
SCR F94
26.80 f0.549
7.16 f0.366
felt
log(M,)
= 27.78 ffl.15
-
_
(M = 7.82 fo.10)
.
- -
V
6.13
SCR F94
27.68 f 0.764
7.75 fO.509
VI
6.04
SCR F94
28.31 50.731
8.17 f0.487
VII
5.44
SCR F94
27.98 f 0.748
VIII
-
-
_
.
-
7.95 f 0.499
- -
BEST WEIGHTED VALUES
best weighted value (with AIII): 27.07 f 0.26
St. Vincent (i) calibration factor: +0.71 f 0.15
7.35 f 0 . 1 7
+0.47 fO.10
best weighted value (without AIII): 27.53 f 0.34
St. Vincent (ii) calibration factor: +0.25 f 0.15
7.66 f 0.23
+0.16 fO.10
(i)
(ii)
(iii) Averaged (i) and (ii) calibration factor: +0.48 (f 0.24)
+0.32 (f0.16)
-
1755 Lisbon earthauake
felt
7.17
SCR F94
29.03 f0.653
8.65 f0.435
I11
6.91
SCR F94
28.60 50.481
8.37 f0.321
IV
6.67
SCR F94
28.77 f0.589
8.48 f0.392
V
6.31
SCR F94
28.52 f0.831
8.31 f0.551
VI
6.00
SCR F94
28.18 f0.716
8.09 f0.477
VII
5.60
SCR F94
28.42 f0.818
8.25 f0.545
VIII
5.24
SCR F94
28.80 f0.845
8.50 f0.564
OPTIONS FOR BEST WEIGHTED VALUES
(iu)
(v)
(ui)
Best weighted value, SCR F94: 28.64 f0.25
Systematic error contrib. (eq. 18, Part IJ):
f 0.10
Max. 1andsIidedistance (upperbound): >29.6
Tsunami magnitude Mt (Abe 1979): 28.95 (f0.23)
a. Lisbon: with (i) calibration factor:
29.34 f 0.50
b. Lisbon: with (ii) calibration factor: 28.88 f 0.50
c. Lisbon: with averaged calibration: 29.12 f 0.59
Lisbon: Final Value & Uncertainty: 29.10 f0.59
[from (vi)c ]
() qualitative uncertainty bounds
-
generate a given M,. For a slip of 12 m and moderate stress
and strain drops, a fault length of -200 km is indicated. This
is a comparable dimension to the axial length of Gorringe
Ridge (Figs lc and 12c). A thrust faulting model for an M 8.7
Lisbon earthquake in which Gorringe Ridge is on the hanging
wall provides a reasonable fit to both the tectonics of the
region and the known strong ground motion and tsunami
effects. The most stringent test for the Lisbon fault model
would be numerical modelling of the Lisbon tsunami, but, to
date, the sparse known tsunami data provide only weak
constraints on the tsunamigenic Lisbon source (e.g. Pedersen
et a/. 1995).
E V I D E N C E F R O M LIQUEFACTION A N D
LANDSLIDING
Liquefaction
Both the New Madrid and Charleston earthquakes produced
widespread liquefaction (Figs 1a,b); some liquefaction was
8.39 f 0 . 1 7
i0.07
>9
8.6
8.86
8.55
8.71
M 8.7
(i0.15)
f0.34
f0.34
f0.39
f 0.39
reported on the Iberian peninsula from the Lisbon event but
will not be assessed here. The preserved liquefaction features
of the New Madrid and Charleston regions have been the
subject of numerous investigations (e.g. Dutton 1889; Fuller
1912; Obermeier 1988, 1989; Amick et a/. 1990). The modified
Mercalli scale first mentions ‘sand and mud ejected in small
amounts’ at MMI VIII. Hence it may be possible to use the
extent and severity of liquefaction to estimate the size of the
source earthquakes essentially independently of the Afel,-AvII
isoseismal regressions.
Youd & Perkins ( 1987) developed the Liquefaction Severity
Index (LSI) scale for assessing the severity of liquefaction in
highly susceptible sediment deposits. The LSI scale has similarities with earthquake intensity scales in that it is based on
secondary phenomena, not specific values of ground motion.
LSI criteria are more quantitative, however, in that LSI values
are assigned for specific ranges in size of liquefaction features
such as fissures, dykes or sand boils (blows). A detailed
description of the scale from LSI 1-100 can be found in Youd
0 1996 RAS, GJI 126, 314-344
Seismic moment assessment of earthquakes-Ill
& Perkins (1987) or Youd, Perkins & Turner (1989). It is
based on data from western US earthquakes and takes the
form
log( LSI) = A + B log(R)+ CM,
(5)
where R is horizontal distance in kilometres from the epicentre
or fault plane, and A, B and C are empirically determined
constants. For the western US, A = -3.49, B= - 1.86, and C =
0.98 (Youd & Perkins 1987);eq. ( 5 ) with these values is plotted
with M as a parameter in Fig. Il(a). It is anticipated that this
relation should not be used without modification to estimate
the size of SCR earthquakes. In particular, 8,the slope of the
M curves in Fig. ll(a), will depend on crustal anelastic attenuation, which differs markedly between West and East (Fig. 4),
and the coefficient C of M will depend on earthquake source
parameters, especially stress drop (Hanks & Johnston 1992).
As examined in Part 11, within a large scatter, SCR earthquakes
do not appear to differ significantly from their active crust
counterparts in average stress drop. This suggests that, to a
good approximation, the M coefficient for SCR North America
should remain C=0.98, the value used for this analysis.
The LSI coefficients are determined empirically. The only
ENA earthquakes with extensively mapped strong liquefaction
are New Madrid and Charleston, the very events for which we
wish to estimate size. Therefore, since an LSI relation for
eastern North America cannot be determined directly, I assume
that, for sites of equal liquefaction susceptibility and for
earthquakes of similar stress drop and M, zones of intense
liquefaction (LSI 100) will be of equal size in the eastern and
western US. The assumption follows from the argument of
Hanks & Johnston (1992) that the pronounced difference in
anelastic attenuation between East and West does not come
into play until epicentral distances exceed 100-150 km. Even
the large and intense liquefaction field of the New Madrid
sequence has an R,,, N 50-60 km, significantly below 100 km.
Moreover, recent work by D. Keefer (personal communication 1995) reporting regional invariance in the attenuation of
liquefaction effects provides additional support for the sizeequality assumption.
The assigned LSI data points of Youd, Perkins & Turner
( 1989) for New Madrid and Charleston are listed in Table 9
and added to Fig. ll(a). The heavy lines are least-squares fits
to the Charleston and New Madrid data. As expected they
exhibit a much slower decrease of LSI with increasing R than
holds for the western US. Moreover, both the New Madrid
and Charleston fits have nearly identical slopes, suggesting
that highly susceptible deposits were abundant and available
in both meisoseismal regions. The LSI-100 effects extend to
-50-75 km from the New Madrid epicentral zone, whereas
equivalent values for Charleston are confined to within
8-10 km. Under the assumptions given above, these R,,,
values should be valid indicators of M based on the western
US LSI formula, eq. (5). The least-squares Rloo value of 8.6 km
yields M 7.4 for Charleston, and the value of 54 km yields M
8.9 for New Madrid.
The Charleston M 7.4 LSI-100 value is consistent with and
provides independent corroboration for the M estimate derived
from isoseismal regressions. The LSI-100 estimate of M 8.9
for New Madrid is much larger than the intensity-based
estimate of M 5 8.1. This should be expected, however, because
the combined intense liquefaction field from at least three
spatially separate major earthquakes would distribute the
-
-
-
0 1996 RAS, G J I 126, 314-344
337
LSI-100 effects to greater distances than would be the case for
a single event. Hence M 8.9 indicates an upper bound for any
single New Madrid event. Therefore, the principal conclusion
of this section is that the liquefaction effects of the Charleston
and New Madrid earthquakes are consistent with-and in the
case of Charleston, support-moment
magnitude estimates
obtained by isoseismal area regression.
Landsliding
Further supporting evidence for the sizes of the three earthquakes analysed above comes from the maximum distance of
triggered landslides, utilizing the work of Keefer (1984). Keefer
analysed a global data set to obtain empirical upper bounds
on the maximum distance from both the triggering event’s
epicentre and its closest point on the fault or rupture plane
for different categories of landslides. The bounds for rockfalls,
a type of disrupted slide or fall (Keefer’s category I), are
reproduced in Fig. 11(b). He developed a more robust empirical upper bound using the total area of induced slides, but
total landslide area could not be estimated for these earthquakes; in fact, hardly any landslide documentation exists at all.
For the Lisbon earthquake, landslide documentation outside
of Portugal consists solely of data points (without descriptions)
in Spain compiled from the 1756 survey of the Royal Academy
of History of Spain (Fig. 7 of Martinez Soares, Lopez Arroyo
& Mezcua 1979). The more remote sites in SCR Spain, shown
in Fig. 10(a),are a minimum of 650-900 km from the Lisbon
source zone, the range reflecting the uncertainty of the Lisbon
epicentre. For New Madrid, the extensive landsliding along
the eastern bluffs of the Mississippi River flood plain (Fig. la)
(Jibson & Keefer 1988) is not useful in terms of maximum
distance of triggered slides. There is a single report (Nelson
1924) of induced rockfalls at Rock Island, Tennessee (Fig. 4),
-400 km from New Madrid, which is currently unverified
(J. Johnston, personal communication 1996); I include it in
Fig. 11(b) for reference only. The single landslide account for
Charleston is from Dutton (1889, p. 331), reporting large
rockfalls in the Black Mountain region of North Carolina (see
Fig. 8), -330 km from the centre of the MMI X Charleston
isoseismal. The coastal plain location of Charleston affords
little opportunity for (Category I ) landsliding, so the absence
of other landslide reports is not surprising.
Fig. l l ( b ) includes the landslide data for the three earthquakes studied as horizontal lines. The Charleston line intersects the Keefer (1984) bounds at M 7.6-8.6; for New Madrid
the range is M 8.0-9.2; and the exceptionally large distances
for Lisbon yield 2 M 9. Almost all of the data used by Keefer
to construct these landslide upper bounds was from active
tectonic regions and has not been calibrated or corrected for
stable continental regions. The effects of the low attenuation
of SCR were minimized for the liquefaction data by considering
only close-in, high LSI values. The opposite applies to the
landslide data where maximum distances are well beyond
100-150 km. Calibration to SCR crust would undoubtedly
lower the M ranges obtained from Fig. 11(b).Such a reduction
would bring the landslide-inferred magnitudes of the earthquakes studied into closer agreement with the isoseismal-based
estimates.
-
338
A . C. Johnston
1
>
c
._
L
a,
>
a,
v)
0
Charleston
10
100
Horizontal distance from fault zone ( R , km)
1
1000
(b) Landslide evidence
.....................
4.0
4.5
5.0
5.5
6.0
6.5
7.0
Magnitude
7.5
8.0
8.5
9.0
9.5
(M)
Figure 11. (a) The liquefaction severity index (LSI) of Youd & Perkins (1987) developed for the western US (dashed lines) LSI data from eastern
North America are superposed The description of LSI data points for New Madrid and Charleston are in Table 9 (from Youd, Perkins & Turner
1989) The least-squares fits to these data are shown by solid lines (b) The Keefer (1984) relationship between maximum distance of landsliding
(disrupted slides or falls) and moment magnitude The heavy lines outlining the shading are Keefer’s approximate upper bounds to the epicentre
and to the nearest point on the fault or rupture zone, respectively. No least-squares fit to the data was attempted so these upper bounds correspond
to the minimum, not best average, magnitudes observed to produce landslides at a given distance. Distances to rockfalls caused by the three study
earthquakes are shown by dashed lines: 330 km, Charleston; -400 km, New Madrid 660-850 km, Lisbon (see text or Figs 4, 8 and 10 for sources).
CONCLUDING R E M A R K S
The tectonic settings of the New Madrid, Charleston and
Lisbon earthquakes differ from each other as well as from the
plate boundary/active crust environment far more typical of
large
- and -great earthquakes. The New Madrid sequence
occurred in a failed intracratonic rift that underwent several
phases of crustal extension and intrusion from the Late
Q 1996 RAS, G J I 126, 314-344
339
Seismic moment assessment of earthquakes-Ill
Table 9. Estimated LSI for the 1811-1812 New Madrid and 1886 Charleston earthquakes (from Youd, Perkins & Turner 1989).
SITE
DESCRIPTION
REFERENCE
DIST. ESTIMATED LSI
km
Max Best Min
NEW MADRID
large fissure, sand boils and dis laced
Fuller 1912
4 0
p u n d from a prox. Marked s e e , Ark.
o New Madrig, Mo.
North of New Madrid, fissures 2-4 ft
Fuller 1912
50-75
wide "the edges bein marked with
p. 54 82-83,
and dg. 1
accumulations of san%and dark clayey
shale." Very large sand boils also
occurred in this area.
Fuller 1912
South of Marked Tree, Ark.,
p. 56, figs. 1 & 7
fissures 10-15 ft deep and10-20
ft wide (most likely grabens)
"Stream banks were caved northward to Fuller 1912
140
Herculaneum, northeastward to Indiana, p. 53
and southward to the Arkansas"
Less severe ground failure henomena Street & Nuttli 1984
p. 50
were observed along the Oko River as
far up river as a oint somewhere between
present day Henlerson and Owensboro, Ky,
and as far south alon the Mississippi .
h v e r as the mouth the St. Francis River.
"O.W. Nuttli (oral comm.) has found
Obermeier 1988
275
p. 12
historical accounts that mention sand
blow activit on the flood plain of the
Mississippi k v e r , near St. Louis"
"Ben-+ (1908) tells of.sand blows
Street & Nuttli
250
1984, p. 45
and !ssuring in Whte County,
Illinois along the Wabash Rwer."
interior of zone of
principal disturbance
boundary of zone of
principal disturbance
bound of zone of
minor disturbance
100 100
LOO
100
80
40
30
50
09
near St. Louis, Mo.
White County, Ill.
10
10
20
10
CHARLESTON
Mile 12 + 700
Northeastern kailroad
(Ten Mile Hill)
"Superstructure on embankment at
Peters &
point where trestle succeeds earthwork
Herrmann 1986
shifted 8 ft 4 inches to w. embankment p. 53
forced along azimuth down p d e
jamming south end of trest e
Ten Mile Hill
Dutton 1889
"These craterlets seemed to reach their
p. 284
greatest develo ment, both in size and
number, near &n-Mile Hill. Some of
them were ve large, one measuring
21 ft across. a a n acres of ound were
overflowed with h e sand, w%ch was tyo
feet or more in thickness near onfices.
8
100 100 100
10
100 100
80
most distant laces of
marked dispyacement
Map showing "localities of marked
honzontal dis lacement."
(point at Charreston disregarded)
Dutton 1989
plate 27
10
100
80
70
most distant places of
of large sand boils
Map showing,"areas conspicuous
for craterlets.
Dutton 1889
plate 28
24
50
30
20
most distant places of
of horizontal displ.
Map showing "localities of
honzontal displacement."
Dutton 1889
plate 27
30
30
15
10
bound of sand boils
noted by Dutton
Map showin$ "line limiting zone
of craterlets.
Dutton 1889
plate 28
35
20
10
5
90
5
3
1
news aper report
"Ne.ar.theresidence of Senator B.H.
Williams was a sink of at least one
Q.6 bermeier,
foot in depth ...Three or four heaps of
SGS, written
ay sand stood at irre ular intervals on commun.)
g e ed es of the sink. h e y were
r b d y thrown up through the fissure
y the internal convulsion.
Georgetown, South
Carolina
Precambrian to the Cretaceous. Charleston's setting was the
thinned and intruded crust of a Mesozoic passive margin. The
Lisbon earthquake ruptured 130 Myr oceanic lithosphere
within a broad compressional deforming zone between the
stable Eurasian and African plates. Current seismotectonics of
-
0 1996 RAS, G J I 126, 314-344
the source regions suggest that Lisbon and NM3, the last New
Madrid major event, were thrust earthquakes, NM1 and NM2
were strike slip, and Charleston-most speculative of all-also
strike slip.
There were two principal objectives of this study. The first
340
A . C . Johnston
(a)
Viable Fault Models
New Madrid
(i) NM 1: Strike-slip. vertical fault, Log (Mo) = 20.25,
M = 0.1
L = 140km
W = 33km
(16 Dec. 1011)
n
ii) NM 2: Strike-slip, vertical, Loa [Mn) = 27.00.
M
= 7.0
I,
t.
F
= 33km
d = 7.5m
c
\
(23 Jan. 1812) W
-
(iii) NM3: Dip-slip, thrust, dip-32’, Loq (M0)=20.00, M=B.O
/ ’
t
1
NM3
(7 Feb. 1812)
“”“J
(b)
’
1
1
1
I
+
L=75km
W = 45km
d = 0.0m
25km
1-
Charleston: Log (M,) = 27.00, M 7.3
strike-slip, vertical fault
(i) High geotherm, hl= 16km
L = 50km
W=16km
d=3.4 m
80 5’W
80 0“
33 00
(ii) Low geotherm, h,= 25km
-
W=25km
3.4 m
a=
32 5”N
Figure 12. Viable fault models for the (a) New Madrid, (b) Charleston and (c) Lisbon earthquakes. All fault models, maps and profiles have the
same distance scale. The illustrated faults are the preferred (boxed) models of Tables 5 and 7. These models are non-unique, but each represents a
and a fault length L that is in accord with the source zone’s
possible mode of rupture with a static strain drop A& that does not exceed 2 x
tectonics as depicted in the map and profile views of each. The seismic moment M , is that computed from the isoseismal regressions of this study.
Faulting depths (fault widths) are limited according to temperature-depth profiles (see text for fuller discussion). Map views are taken from Fig. 1,
except that source faults are added, the Blytheville arch is added for New Madrid, and the bathymetry for Lisbon is in fathoms (1 fa= 1.83 m).
See Johnston & Schweig (1996) for the development of the New Madrid faulting model.
was to determine the seismic moments of the New Madrid,
Lisbon and Charleston earthquakes with specified uncertainties; the second was to explore viable source parameters and
faulting models given their tectonic settings and estimated
sizes. The first objective was accomplished using regression
analysis on the known distribution of intensities, supplemented
by evidence from liquefaction, landsliding, and tsunami runup. Calibration procedures were necessary for all earthquakes
studied: for New Madrid and Charleston because east-central
North America has demonstrably lower anelastic attenuation
0 1996 RAS, GJI 126, 314-344
Seismic moment assessment of earthquakes-Ill
(c) Lisbon: Log (Mo) = 29.10, M 0.7
dip-slip thrust fault, dip -40"
.fault base @ 600°C isotherm (hz50km)
0
t
L = 200 km
W= 00km
iT=
12m
Oceanic Lithosphere
p = 6.5 x 10" dyne/crnZ
-
p = 3.3 gm/cm3
D
/ Gorringe
Ridge
Figure 12. (Continued.)
Table 10. Summary of size estimates: New Madrid, Charleston and Lisbon.
New Madrid
Mo= 1.8 (0.60-5.3) x 10'' dyne-cm
(0215) 16 Dec. 1811
23 Jan. 1812
New Madrid:
M0=6.3 (2.0-20.0) x loz7 dyne-cm
07 Feb. 1812
New Madrid
M,= 1.0 (0.32-3.1) x lo2*dyne-cm
New Madrid
M0=6.3 (2.2-18.2) x
dyne-cm
(0815) 16 Dec. 1811
Charleston:
M,= 1.0 (0.42-2.4) x lo2' dyne-cm
31 Aug. 1886
Mo= 1.3 (0.32-4.9) x loz9dyne-cm
01 Nov. 1755
Lisbon:
0 1996 RAS, GJI 126, 314-344
M = 8.1 k0.3 1
M=7.8+0.33
M = 8.0F0.33
M =7.2+0.31
M = 1.3 0.26
M = 8.7 k0.39
341
342
A. C . Johnston
than average global SCR crust; again for New Madrid because
its western isoseismal areas were unknown; and for Lisbon
because a large percentage of its isoseismal areas consisted of
oceanic lithosphere or active tectonic continental crust. Hence
seismic moment estimates of the study events were obtained,
but not without complications that increased the associated
uncertainties. The final determined moments and uncertainties
are summarized in Table 10.
The uncertainties are symmetrical about the best average
values. However, from the weight of the supporting liquefaction, landslide and tsunami data and the conservative application of the intensity data, I believe that the possibility is
greater that the final quoted moments could underestimate
rather than overestimate the true moments of these earthquakes. That is, using all the data at hand, it would be easier
to argue for seismic moments that are larger rather than
smaller than the best-weighted averages given here.
I approached the second objective by considering the geophysical setting of each of the earthquakes. This provided
primary constraints on crustal/lithospheric rigidity and the
permissible depth of faulting, which were then combined with
the above M , values to generate a range of permissible-not
necessarily physically plausible-faulting models. The ranges
in Tables 5 and 7 show the more plausible, or viable, models
in bold type. From these I have selected one or two models
for the summary Fig. 12. This figure is not intended to show
the actual fault ruptures of the study earthquakes; rather the
intent is to depict what their source faults could be. Dimensions,
slip and rigidity must yield the M , determined from regression.
Stress and strain drops should not be overly extreme. Depth
of faulting should not violate the estimated limits imposed by
geothermal and rheological considerations. Scenarios for the
New Madrid sequence are the most complicated and depend
strongly on the synoptic shear-zone model for continental
crust (Scholz 1988, 1990).
The final results of this study are best stated in general
terms even though specific values and uncertainties are available above. I conclude that two of the three principal
1811-1812 New Madrid earthquakes had magnitudes in the
low M 8 range; the other was in the high M 7 range. NM1
was the largest, but, with the associated uncertainties, their
sizes are statistically indistinguishable, clustered around M 8.0.
The largest aftershock was in the low M 7 range. Coseismic
ruptures could have extended to lengths in the 60-180 km
range and to depths significantly deeper than indicated by
current seismicity. Displacements probably were in the 6-11 m
range, yielding static stress drops in the 40-90 bar range. The
magnitude of the Charleston earthquake was in the low- to
mid-M 7 range. Its most probable fault length was in the
30-80 km range; fault width, 16-25 km; and fault slip, 2-4 m,
yielding a static stress drop of about 50 bar.
Finally, I find that Lisbon was a truly great earthquake with
magnitude in the upper M 8 range. This requires a large fault
of the order of 180-280 km length and 50-80 km width, a
large average displacement probably exceeding 10 m, a high
lithospheric rigidity, and an average 40- 100 bar static stress
drop. Because of the compressional tectonics offshore from
Portugal and the great tsunami generated, a thrust faulting
mechanism can be assigned to Lisbon with relatively high
confidence. The 1755 Lisbon earthquake was probably a major
contributing episode in the continuing tectonic uplift of
Gorringe Ridge.
ACKNOWLEDGMENTS
This paper was critically reviewed by J. Dewey, T. Hanks, N.
N. Ambraseys, G. Bollinger, M. Ellis, L. Youd, and associate
editor S. Ward; I thank them all for significant manuscript
improvements. At a fundamental level, historical rather than
seismological data govern its conclusions, so I wish to acknowledge those whose decades of work has preserved and interpreted the known information about these important
earthquakes, especially Ron Street and the late Otto Nuttli for
New Madrid and Gil Bollinger for Charleston. I thank Shamita
Das for stimulating my interest in the Lisbon earthquake even
though it is not an SCR event. The critical interest of
K. Coppersmith, T. Hanks, G . Bollinger, R. Wheeler, E.
Schweig and J. Gomberg prompted me to pay much greater
attention to specifying uncertainties. I thank K. Satake,
D. Keefer, B. Mitchell, C. Scholz and A. Bent for information
or preprints on tsunamis, landslides, Q, fault slip, and ENA
earthquakes, respectively, J. Johnston for historical research
and T. Broadbent for drafting. Some phases of this work were
supported by the Electric Power Research Institute and the
Tennessee Centers of Excellence program. Contribution
number 268 of the Center for Earthquake Research and
Information, the University of Memphis.
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