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Australia: The Land Where Time Began |
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Indian Ocean Dipole – Relation of Deep Meridional Overturning Circulation in
the Indian Ocean
For this paper the 42-year German Estimating and circulation and Climate of
the Ocean (GECCO) synthesis data to analyse and examine the relation of the
deep meridional overturning circulation (DMOC) with the mode of the Indian
Ocean Dipole (IOD). Decomposing the DMOC into the Ekman and geostrophic
transport, the external mode and a residual term are used to assess the
contrition of various dynamical processes. The DMOC with a marginal residual
term is successfully described by the first 3 terms. The following
conclusions are obtained:
1.
The seasonal cycle of the DMOC is determined mainly by the Ekman component.
During the transitional seasons (March-April and September-October) in the
northern Indian Ocean Basin, where the geostrophic component dominates.
2.
The Ekman component dominates the structure of the DMOC at the beginning of
the IOD phase (May-June); at and following the peak phase of the IOD
(September-December, the structure of the DMOC is determined primarily by
the geostrophic component in correspondence with the sea surface temperature
anomalies which are well developed, while a secondary role is played by the
wind, and therefore the Ekman component, south of 10oS and
contributes negatively within the zonal band of 10o on both sides
of the equator. Therefore, there is surface to a deep ocean connection
through which wind and ocean temperature anomalies that are related to the
IOD are transferred down to the deep ocean. Even in the deep ocean signals
that propagate westwards are observed, which suggests possible roles for the
Rossby waves in the transferring of the surface signal to the deep ocean.
Introduction
A deep meridional inflow is one of the characteristic features at the
southern Indian Ocean where North Atlantic Deep Water, Antarctic Bottom
Water, and Circumpolar Deep Water enter the Indian Ocean in the western
basin off Madagascar, as well as off the coast of East Africa and in the
east, along the ninety East Ridge (Mantyla & Reed, 1995; Schott & McCreary,
2001). The Indian Ocean deep meridional overturning circulation (DMOC) and
its temporal and spatial variability remain open to be investigated in a
manner that is more systematic.
Based on individual transoceanic hydrographic sections, either estimating
the geostrophic transports (Toole & Warren, 1993; Robinson and Toole, 1997)
or by an inverse box modelling approach (Macdonald, 1998; Ganachaud et
al., 2000; Bryden & Beal, 2001;
Sloyan & Rintoul, 2001; Ganachaud, 2003; Lumpkin & Speer, 2007; McDonagh et
al., 2008), the time-mean
structure of the DMOC was obtained earlier. The results of these earlier
studies are often contaminated by strong seasonal variability in the Indian
Ocean, as both earlier approaches depend heavily of hydrographic data
sections, which are often sparse in the spatial and temporal resolution. In
comparison, Numerical model efforts as an alternative way show weak, though
noticeable, time-mean bottom inflow [2-5 Sv (1 Sv ≡ 106 m3
s-1); Garternicht & Schott, 1997; Lee & Marotzke, 1998;
Stammer et al., 2012; Drijfhout &
Garabato, 2008: Wang et al.,
2012). The estimation of 17 Sv from Ferron & Marotzke (2003) is the only
exception, which has been shown to not be maintainable when an adjoint
method modifies the initial conditions and forcing fields that are applied
for several hundred years (Palmer et
al., 2007) instead of for a few years as was the case of Ferron &
Marotzke (2003). It has also been acknowledged, on the other hand, that
although the results of numerical modelling are dynamically consistent they
are conditioned to prescribe weak vertical mixing, different surface heat
and momentum fluxes products, model resolution, boundary conditions that are
open, and deficits with respect to the formation of bottom water (Schott &
McCreary, 2001; Drijfhout & Garabato, 2008; Rintoul et
al., 2010; Wang et
al., 2010).
Large seasonal variations over most latitudes are displayed by the DMOC and
the DMOC is characterised by 2 counterrotating meridional overturning cells
(Lee & Marotzke, 1998; Wang et al.,
2012): In the boreal winter one cell is anticlockwise (facing the paper (or
monitor) unless specified otherwise), located in the region couth of 20oS,
while the other is clockwise, and is located in the interior of the Indian
Ocean centred on 10oS. Responding to the seasonal reversal of the
monsoon both cells reverse their rotation in the boreal summer. Almost the
entire water column is covered year-round by the overturning streamfunction,
which indicates the deep-reaching effect of the Indian monsoon (Schott et
al., 2002).
Pronounced interannual variability is exhibited by the Indian Ocean DMOC. It
has been found that there is a significant 2-year variation in the DMOC
(Wang et al., 2012) as well as
other oceanic variables that are essential, such as meridional ocean
transport of heat in the Equatorial Indian Ocean (Chirokova & Webster, 2006)
which is reminiscent of the tropospheric biennial oscillation (TBO; Meehl &
Arblaster, 2002; Loschnigg et al.,
2003). There is high correlation between the DMOC and the most prominent
interannual mode, the Indian Ocean Dipole (IOD) (Wang et
al., 2012), the positive phase of
which is characterised by negative SST anomalies in the tropical
southeastern Indian Ocean and positive anomalies in the tropical western
Indian Ocean.
The characteristic dipole SST pattern of IOD events first occurs in
May-June, peaking in September-October; its demise is likely to be the
reversal of the monsoon that weakens winds along the equator and along the
coast of Indonesia, which diminishes the importance of oceanic dynamics in
the regulation of SST (Saji et al.,
1999). It is suggested the acknowledgement that the IOD is parented by an
anomalous monsoon through the generation of zonal temperature gradients
between upwelling regions, it was suggested by Webster et
al. (1999, 2002) that the
east-west SST gradient that were initiated by the anomalous monsoon
intensity can lead to the enhancements of the zonal gradients of the SST by
modulating local upwelling and meridional transport of oceanic heat through
involving equatorial wave dynamics. In this respect, they suggest that the
IOD is an integral part of the TBO, as it could regulate the intensity of
the monsoon system, introduces slow dynamics into the SST anomalies that are
initiated by the monsoon, and help to penetrate these anomalies to the next
year.
In this study Wang et
al., show that the Indian Ocean DMOC is linked to conditions of the
ocean surface mainly through wind (thus the Ekman component) on seasonal
scales and through anomalies of wind and sea temperature on interannual time
scales that are related to the IOD, by decomposing the DMOC into 3 dynamical
components and assessing their relative contributions. They extended the
traditional view of the air-sea coupled system in the Indian Ocean, which is
confined to only the upper Indian Ocean and the atmosphere overlying it so
far (Webster et al., 2002; Feng &
Meyers, 2003; Loschnigg et al.,
2003; Annamalai &Murtugudde, 2004; Pillai, 2008). The organisation of this
paper is:
·
section2 described data and methodology;
·
section 3 presents the roles of various dynamical components of the DMOC
during IOD events;
·
section 4 is a discussion;
·
section 5 provides a summary and conclusions.
DMOC variations during IOD events
A high-pass Butterworth filter with a cutoff frequency of 7 years was
applied to the meridional streamfunction, SST and the wind stress field at
each model grid for the period 1960-2001, in order to function on
interannual variations. Of these variables, components that were longer than
7 years which have been removed through the high-pass filtering, are
presented in fig.3, and the remainder of the analysis (except for fig. 8)
revers to the variables that are high-pass filtered. Resulting from this,
there 2 dominant interannual periodicities and well as addition to seasonal
cycles, are retained in their analysis: 2-3 years which correspond to TBO
and 3-7 years which correspond to ENSO (Pillai & Mohankumar, 2009). Note,
that prior to the composite analysis, the high-pass filtered variables are
further subject to deseasonalisation; i.e., the monthly climatological
states are removed from the respective variable.
Climatological states: DMOC and surface wind
Wang et al. first presented the
annual cycle of the Indian Ocean deep meridional streamfunctions that have
been high-pass filtered.
In the boreal winter (from November to February) and summer (from
May-August), meridional overturning cells are most obvious double
counterrotating, while in transitional seasons (March-April and
September-October) the northern cell splits into 2 parts. A clockwise closed
deep cell dominates the interior of the ocean (north of 20oS) in
the boreal winter and an anticlockwise deep cell locates in the southern
Indian Ocean (south of 20oS), both of which reverse their
rotations in the boreal summer, responding to the reversal of the summer
monsoon. The entire water column is covered by this pair of deep cells with
maximum strength of 10-12 Su at a depth of 1,000 metres (located between 10oS
and 5oN) and is separated noticeably at around 20oS
throughout the year.
In the boreal winter (November-February) positive wind stress curl occurs
north of 10oS and to its south negative wind stress curl occurs
with the maximum centred along 20oS. In the Boreal summer the
wind stress pattern resembles that in winter months, though with opposite
signs. According to Ekman dynamics, positive (negative) maximum curl
corresponds to maximum downwelling (upwelling in the Southern Hemisphere, ha
has been manifested by confluences of upwards (downwards) branches of both
cells near 20oS.
The low-frequency wind component (≥7 years ), which is primarily
characterised by southeasterlies in the southern Indian Ocean and over the
circumpolar ocean, strong westerlies, persists throughout the year and Wang
et al suggest it may relate to the
southern annular mode (Thompson & Wallace, 2000). The Ekman drift that
results leads to convergence and downwelling between 60oS and 20oS
where the maximum easterly component is located, and between the equator and
20oS divergence and upwelling, which accounts for an upwelling
(positive) cell within the southern Indian Ocean domain with maximum
amplitude of about 2 Sv in the deep ocean.
Climatological states: DMOC and its dynamical components
The main structure of the DMOC is determined by the Ekman components and
account for a large part of the amplitude throughout all months. In
transitional seasons (March-April and September-October) there are
exceptions in the northern Indian Ocean Basin, where the Ekman contribution
is offset by the geostrophic component which dominates the local structure
of the DMOC. There are 2 main features that are shown by the external
component:
1)
South of 20oS there is a coherent pastern throughout the year,
which extends to 5,000 m, and is in phase with the Ekman contributions for
most of the year.
2)
In summer (May-October) and winter months (November-April) its spatial
features are rather similar though with opposite signs.
Because of the zonal topographic gradients, changes in horizontal ocean
gyre, or in cross-basin zonal gradients of the meridional velocity, which is
depth independent, in relation to the summer/winter monsoon shift, both
features reflect only.
In summary, the structure of the seasonal DMOC can be reconstructed
satisfactorily from its 3 dynamical components: Ekman flow and its
barotropic compensation, the thermal wind component, and it external mode
which is depth independent. In the following, relative contribution s of
these 3 dynamical components to DMOC are assessed each month and are subject
to a composite analysis in order to obtain a quantitative estimate of their
roles in relation to the IOD events.
Positive IOD composites
Composite analysis for strong positive IOD events is used to study possible
imprints on the Indian Ocean DMOC of IOD events. First, these selected
events are subject to deseasonalisation: i.e., climatologically monthly DMOC
is removed prior to the composite analysis.
The strength of the IOD is measured by differences in zonal sea surface
temperature (SST) between the tropical western Indian Ocean (50oE-70oE),
(10oS-10oN) and the tropical southeastern Indian Ocean
(90oE-110oE, 10oS-0o), following
the same definition as is given in Saji et
al., 1999). As is shown in fig. 5
in this paper, the time series of the IOD index along with that of the
Indian Ocean DMOC index is represented by the volume of transport of the
bottom flow (ƴ1 ≥ 28.11 kg/m3) across 34oS.
A 7-month running average is used to smooth both time series. There were 90
positive IOD that were obtained by Wang et
al. (1965, 1967, 1972, 1976, 1977,
1982, 1987, 1994, and 1997), with the threshold of 1.5 (multiplied by the
standard deviation). The DMOC index is out of phase with the IOD index, in
most cases, whereas the years 1065 and 1987 are exceptional. The years 1965
and 1987 were excluded in order to highlight the commonality of the majority
cases, thereby keeping 7 positive IOD cases in the following composite
calculations.
Off the coast of Java cold SST anomalies first appear in May, accompanied by
stronger than normal alongshore southeasterlies and easterly anomalies along
the equator between 80oE and 100oE. 2-fold effects of
these wind anomalies have been revealed by previous studies: stronger
coastal upwelling off Java and Sumatra is led to by alongshore southeasterly
wind, which is stronger than normal, while along the equator the easterly
anomalies hamper the eastwards intrusion of the equatorial current, thereby
reducing the supply of heat to the east. Cooling off Indonesia is enhanced
by both effects, where another prerequisite for this cooling is provided the
shoaling thermocline (Xie et al.,
2002). This process may relate to coastal advection as well as equatorial
and Kelvin wave dynamics (e.g., Feng & Meyers, 2003). E.g., (Xie et
al., 2002) that a westwards downwelling Rossby wave was forced by
the wind stress curl that was associated with the IOD leads to deeper
thermocline and therefore and the warmer SST in the west, which is
intensified further on the surface as a result of reduced wind speed and
evaporation that is induced by anomalously extended northwestwards–extended
southeasterly trade winds (Saji et al.,
1999). Therefore, changes of the southeast trade winds on the equatorial
region coupled with SST anomalies through ocean dynamics give rise to SST
dipole events. The westwards propagation of cold SST anomalies along 10oS
from September to December is a partial manifestation of the major role
played by equatorial wave dynamics in the regulation of the dipole SST
dynamics. The IOD-related atmospheric circulation, on the other hand, is
essentially a reflection of the zonal pressure difference in the equatorial
Indian Ocean; it is thereby it is related closely to changes in equatorial
wind systems, such as the Walker circulation (Saji & Yamagata, 2003).
Northwesterly wind anomalies appear in the central southern Indian Ocean in
September-October, which indicates a weakening of the southeasterly trade
winds. They form, together with the easterly wind anomalies along the
equator, an anticlockwise wind stress curl that generates anomalous Ekman
pumping within the zonal band from the equator to 200S. In
November-December this positive wind stress curl extends further westwards
along the equator, and occupies almost the entire southern tropical Indian
Ocean.
The DMOC composite is characterised in May-June by a dipole demarcated at
about 10oS; In July-August the associated anomalies almost
vanish, exhibiting no distinguishable spatial patterns. An anomalous
negative cell occurs between 20oS and the equator in
September-October, which expands into November-December into the entire
southern Indian Ocean Basin accompanied by a positive cell in the upper
2,000 m in the northern Indian Ocean.
The DMOC features that are IOD-related shown above are, to a large extent,
described by the reconstruction following Eq. 1 (in Figs. 7a,e,f) in this
paper by Wang et al., 2014). At
different phases of the IOD each dynamical component of the DMOC plays
varying roles. In May-June, e.g., the DMOC anomalies are mainly contributed
by the Ekman components which are characterised by a pair of cells that are
reversely rotating separated at 10oS, though in later months it
is geostrophic current that determines the spatial feature of the DMOC
anomalies; in particular, stratified DMOC structures in November-December in
the Southern Ocean basin is attributed mainly to the geostrophic component.
Ekman-related DMOC anomalies are offset by wind contribution on both sides
of the equator within 10o of the equator by thermal wind
contribution, though at south of 100S, the Ekman thermal wind
components share the same sign; therefore both contribute to the strong DMOC
anomalies in the southern Indian Ocean. The DMOC anomaly that is related to
the external mode is confined primarily to the deep southern Indian Ocean
below 1,000 m and its highest loading in September-December is near 10oS.
The
IOD signals that are the most pronounced are manifested as wind and SST
anomalies. These 2 variables link directly to the Ekman and thermal wind
components of the DMOC, respectively, through which the IOD signals
penetrate to the deep ocean. The SST anomalies are still weak at the
starting phase of the IOD events, May-June, and so are the anomalous wind
flows. The associated anomalous wind stress curl on the other side exhibits
cross-basin coherent patterns: anomalous bands of positive wind stress curl
anomalies along approximately 10oN and 20oS, which
results in the local DMOC extrema. The wind stress curl anomaly loses its
cross-basin coherent feature in July-August; correspondingly, the DMOC
contribution that is Ekman-related reduces considerably. Meanwhile, near the
Java coast the cold SST anomaly spreads and enhances further until December,
which is consistent with the strong contribution from the thermal wind
component of the DMOC.
Accompanying strengthening and the propagation to the west of the cold SST
anomalies in the southern tropical Indian Ocean from September to December,
the anomalous equatorial easterly and northwesterly winds in the southern
Indian Ocean strengthen and, moreover, resume a cross-basin coherent feature
with the zero (anomalous) wind stress curl line taking up a
northwest-southeast orientation between 10oS and 20oS.
This corresponds with the negative Ekman component of the DMOC that is
centred along around 15oS.
Discussion
In
this paper Wang et al. adopted the
method of Lee & Marotzke (1998) and Hirschi & Marotzke (2007) and applied it
to the DMOC in the Indian Ocean on seasonal and IOD-related timescales. Wang
et al. also analysed transitional
months when the Asian monsoon reverses its direction (March-April and
(September-October), and the DMOC features a 2-cell structure to the north
of 20oS in contrast to a coherent structure during
monsoon-prevailing months as a supplement to Lee & Marotzke (1998) that
discussed the DMOC feature in the monsoon-prevailing months (January and
July). The previous prevailing monsoon weakens in the previous months and
the subsequent monsoon regime is not yet fully established; a result from
this is that the Ekman component weakens, which is seen more clearly to the
north of the equator, whereas a coherent structure is obtained by the
geostrophic component across the entire Indian Ocean Basin. The latter
partly offsets the former and in March-April even supersedes it to the north
of 10oS and north of the equator in September-October, which
contributes to the 2-cell feature north of 20oS in the
transitional months.
At
around 20oS there is a clear demarcation line which separates 2
circulation cells that are rotating reversely throughout the year. This
reflects the effects of the monsoonal winds and intertropical convergence
zone which is associated with it Southern Ocean. In January this demarcation
line is less pronounced and in July, is almost indistinguishable (Lee &
Marotzke, 1998). According top Wang et
al. the difference may mainly from the difference of wind product: GECCO
datasets used NCEP wind stresses and Lee & Marotzke (1998) used surface
fluxes from Hellerman & Rosenstein (1983). It may be partially due to the
strong constraints in the sponge layer in connection to the Antarctic
Circumpolar Current (ACC) (Lee & Marotzke, 1998). Interannual variability of
the ACC, as well as its interaction with the internal Indian Ocean Basin
which, according to Wang et al. may be nontrivial in shaping the circulation of the feature in
the southern Indian Ocean, are inhibited by these strong constraints. It is
demonstrated by the composite analysis for the positive IOD events that the
link between the Indian Ocean DMOC and surface conditions of the ocean:
namely, the characteristically IOD-related and SST anomalies are linked
directly to the structure of the DMOC on Interannual timescales.
Wang
et al. extended by this means the
view of the air-sea coupled system in the Indian Ocean which is merely
confined to the upper Indian Ocean and the overlying atmosphere so far
(Webster et al., 2002; Feng &
Myers, 2003; Loschnigg et al.,
2003; Annamalai & Murtugudde, 2004; Pillai, 2008), to the deep ocean. As a
result of such a connection between the surface and the deep ocean events on
the surface of the ocean in the form of anomalies of wind and SST anomalies
are linked dynamically to the variability of the DMOC; it is therefore not
surprising that DMOC is observed exhibits concentrated energy on timescales
that are characteristic of the surface air-sea coupled system, such as on
the typical time scales of TBO an d ENSO (Wang et
al., 2012). The strong
contribution of the Ekman component to the variability of the Indian Ocean
DMOC that is revealed through the year supports the suggestion that wind
forcing on a large scale has a significant role in sustaining overturning in
the deep ocean by way of mixing processes (Guan & Huang, 2008; Huang, 1999;
Nycander et al., 2007; Huussen et
al., 2012).
Through which mechanism do temperature anomalies related to the IOD
influence the deep ocean down to 5,000 m, is a question that naturally
arises, particularly from September to December. Temperature and Meridional
velocity fields are clearly observed to propagate to the west, even at a
depth of 3,000 m, as is shown in a snapshot of temperature anomalies at
2,000 m at 20oS for the period 1971-1973. When approaching the
equator this propagation to the west is more rapid. Wang et
al., conjectured, therefore, that
the surface signal that is related to the IOD may be carried down in the
form of the annual Rossby waves that are driven by the wind which force the
deep ocean by vertical by movements of the thermocline because of
convergence and divergence in the upper ocean which is wind-driven (Kessler
& McCreary, 1993; Johnson, 2011). A strong presence of annual wind-driven
Rossby waves has, in fact, been observed in hydrographic profiles of the
Arabian Sea (Brandt et al., 2002)
and in the equatorial Pacific (Kessler & McCreary, 1993) and is reported to
be coupled with the overlying atmosphere on annual and Interannual scales
(White, 2000a, 2001). It is most likely that these wind-driven Rossby waves
are the messenger responsible for communicating the surface variations down
to the deep ocean. According to Wang et
al. a detailed analysis related to
this aspect is planned for future studies.
Wang
et al. say they would like to
stress that the choice of positive events might vary, depending of the
dataset and the definition used (e.g. Saji et
al., 1999; Saji & Yamagata, 2003;
Du et al., 2013). The analysis
that was presented in their paper was based on the GECCO ocean reanalysis
dataset. Wang et al. missed the
IOD events in 1961 and 1963, compared to previous studies (e.g. Saji et
al., 1999; Saji & Yamagata, 2003;
Du et al., 2013). They conjecture
that the mistargeting in 1961 and 1963 may relate to the slow adjustment to
the initial conditions in the assimilation process in the GECCO data. They
attempted to minimise this effect by disregarding the first 8 simulation
years (1952-1958), which may not be enough, which can be seen by the abrupt
increase in the DMOC index from 1960 to 1063.
Also, a similar analysis has been performed for negative IOD cases. 6
negative events are selected based on the threshold of -1.5: 1964, 1970,
1973, 1975, 1996 and 1998. It is shown by a closer look at the annual course
of the IOD index that positive and negative IOD events are not symmetric
around zero. Though all of the positive IOD events that were selected mature
in autumn, 3 out of the 6 IOD events (1964, 1970 and 1975) reach the minimum
between December and March of the following year and in 1973 and 1998 the
minimum IOD index appears near September. Many studies have already noticed
this symmetry between the opposite phases of the IOD (Ummenhofer et
al., 2013). The wide spread of the
development course of negative IOD events makes the interpretation of the
time-based component less meaningful, though the DMOX index shows an
out-of-phase relation with the IOD index during most of the negative IOD
events. The results presented here may, therefore, pertain only to positive
phases.
Summary and conclusions
Wang
et al. have examined the
variations of the Indian Ocean deep meridional overturning circulation
(DMOC) in relation to IOD events by use of GECCO synthesis. In order to
assess the relative importance of various dynamical processes in causing the
seasonal cycle of the DMOC and in imprinting influences of the IOD down into
the deep ocean, the DMOC is decomposed into 4 components (following Hirschi
& Marotzke, 2007):
1)
The Ekman flow, a
barotropic compensation,
2)
Vertical sheared
flows comprising the thermal wind,
3)
the external mode
that accounts for zonally varying topography, and
4)
a residual term that
includes a geostrophic flow that is associated with friction and nonlinear
effects.
On
seasonal to Interannual time scales the residual term remains relatively
small, and was therefore not discussed in detail in this study.
The
overall seasonal structure of the DMOC is determined by the Ekman component
and also contributes to a substantial part of the amplitude throughout all
months. In the transitional seasons (March to April and September to
October), there are exceptions, when over the entire Indian Ocean Basin the
thermal wind component has the same sign, with an amplitude contribution
that is comparable to the Ekman component. The thermal wind component
offsets the Ekman contribution and determines the local DMOC structure, in
the northern Indian Ocean Basin. Moderate seasonal variations at latitudes
where the Somali Current and the corresponding gyre reverse their directions
on a seasonal basis are caused by the external mode (Lee & Marotzke, 1998).
South of 20oS it shares the same sign with the corresponding
Ekman component, which suggests its close link to the seasonally reversed
wind pattern on both sides of 20oS, and therefore also to the
related Ekman drift.
It
is revealed by the composite analysis of the 3 dynamical components of the
DMOC that each component plays a different role during different phases of
the IOD. The DMOC structure related to the IOD is characterised in May-June
by a dipole pattern that is separated at about 10oS, which is
contributed mainly by the Ekman component, whereas the contribution of the
geostrophic component increases considerably and determines the main feature
of the DMOC. These changes in the relative contribution of the Ekman and
thermal wind components of the DMOC relate closely to the surface features
at different phases of the IOD: namely the SST
anomalies related to the IOD and the associated wind anomalies. Wang, W., et al. (2014). "Deep Meridional Overturning Circulation in the Indian Ocean
and Its Relation to Indian Ocean Dipole." Journal of Climate
27(12): 4508-4520. |
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Author: M.H.Monroe Email: admin@austhrutime.com Sources & Further reading |