Magma Production Rates
for Intraoceanic Arcs
Brian R. Jicha1 and Oliver Jagoutz2
1811-5209/15/0011-0105$2.50
I
DOI: 10.2113/gselements.11.2.105
ntraoceanic volcanic arcs have long been recognized as sites where continental crust is created. Yet, despite their importance to understanding
magmatic systems and the evolution of our planet, very little is known
about their long-term rates of magma production and crust formation.
Constraining both crustal construction and destruction processes at intraoceanic arcs allows for improved estimates of magma production. Our revised
magma production rates for active intraoceanic arcs are consistent with
those calculated for mid-ocean ridge segments that have slow to moderate
spreading rates. This is surprising because magma production at intraoceanic arcs has traditionally been assumed to be significantly less than that at
mid-ocean ridges.
form the oceanic crust contributes
~9% (~4 terawatts) to the total
global heat loss of 42–46 terawatts
(the other 91% of the Earth’s
heat loss is thought to occur via
conduction through the crust and
via mantle convection; Nakagawa
and Tackley 2012). However, these
estimates generally ignore the role
of heat loss related to subduction
zone magmatism. Magma production along subduction zones was
assumed to be about an order of
magnitude less voluminous than
KEYWORDS : intraoceanic arc, magma production rates, delamination
that at mid-ocean ridges (Reymer
and Schubert 1984; White et al.
INTRODUCTION
2006), therefore contributing only a negligible amount to
Most magmatism observed on Earth is located at mid-ocean
the global heat-flow budget (about ~1%).
ridges and above subduction zones3. At mid-ocean ridges,
Here, we critically evaluate the above assumption by
oceanic plates diverge and the Earth’s mantle moves
reviewing recent estimates of magma production and
upward. The reduction in the overlying pressure enables
focusing on intraoceanic subduction zones where volcanic
mantle melting (a process called “decompression melting”)
arcs form on top of older oceanic crust. The rationale for
and leads to magma production and the formation of
focusing on intraoceanic arcs is two-fold. First, continental
new oceanic crust. At subduction zones, water and other
volatile components stored in the subducting oceanic crust arcs have an additional level of complexity compared to
intraoceanic arcs in that they are constructed on preexare released at depth due to the increase in pressure and
isting continental crust, thereby making it challenging to
temperature. This water-rich component fluxes the Earth’s
decipher what fraction of the crust is composed of old versus
mantle, thereby lowering its melting point, a process
new material. Second, continental arcs show pronounced
comparable to salt lowering the melting point of snow and
cyclicity: most of the arc crust is built in several short
ice. This induces melting (referred to as “flux melting”) and
episodes, each of which span 5–20 My. These episodic
produces magma, some of which is erupted at continental
and intraoceanic arc volcanoes, but most of the magma magmatic “flare-ups” have estimated crust production rates
of ~85 km3 km−1 My−1, and are separated by episodes of low
never reaches the surface. Instead, it freezes (crystallizes)
within the crust to form so-called plutonic rocks. Plutonic magma-production rates of ~20 km3 km−1 My−1, which can
and volcanic rocks formed above subduction zones often last approximately 25–50 My (e.g. Ducea et al. 2015 this
issue). The general assumption is that a flare-up is indicahave chemical characteristics similar to continental crust,
and it is a widely held belief that continental crust has tive of a period of unusually high magmatic activity, but
it is equally possible that the quiescent period is unusubeen formed by this process throughout Earth’s history.
ally low. Therefore, in order to interpret the periodicity
Understanding the rates of global magma generation and
observed in continental arcs, it is essential to quantify
crust formation is critical for constraining the evolution
the “normal” background magma production rate. This is
of Earth’s crust and mantle and for quantifying the overall
best done in intraoceanic arcs where periodicity appears
heat-flow budget of the Earth. For example, magmatic heat
to be limited.
loss due to the formation and crystallization of melts that
We will fi rst critically review the existing data on magma/
crust production and propose ways forward to better
constrain it. We demonstrate that magma production
1 Department of Geoscience, University of Wisconsin–Madison
rates at intraoceanic arcs are comparable to those along
Madison, WI 53706, USA
mid-ocean ridges and that the high-magma-production
E-mail: [email protected]
periods observed in continental arcs represent the “normal”
2 Department of Earth, Atmospheric, and Planetary Sciences
arc stage. These results have profound implications for
Massachusetts Institute of Technology
Cambridge, MA 02139, USA
understanding Earth’s heat loss.
E-mail: [email protected]
3 A third location of magmatism is intraplate volcanism
(such as Hawaii), which is not discussed in this paper.
E LEMENTS , V OL . 11,
PP.
105–112
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Schematic
depicting a
mature intraoceanic arc
with processes that add or
remove material from the
arc crust. MODIFIED FROM
FIGURE 5 OF JAGOUTZ AND
SCHMIDT (2013)
FIGURE 1
FLAWS IN EXISTING ESTIMATES
OF MAGMA PRODUCTION RATES
For mid-ocean ridges, calculating the crust production rate4
is relatively straightforward. The average crust production
at mid-ocean ridges is ~18 km3 per year and the global ridge
system is 67,338 km long (Bird 2003), which translates to an
average crust production rate of ~233–326 km3 km−1 My−1,
assuming a 5–7 km thick oceanic crust.
Several attempts have been made to calculate crust production rates at intraoceanic arcs by estimating a date of
arc inception and using seismic profi les to approximate
arc volumes (TABLE 1; Crisp 1984; Reymer and Schubert
1984; White et al. 2006). These estimates are generally
crust production/accumulation rates and not, strictly
speaking, magma production rates. These terms should
not be interchanged because crust production rates do
not take into account the volume of crust that has been
removed from the arc since its formation. Early studies by
Reymer and Schubert (1984) and Crisp (1984) produced
broadly consistent crust production rates for intraoceanic
arcs: ~20–40 km3 km−1 My−1. Dimalanta et al. (2002)
combined gravity data with seismic data to better estimate
the arc volumes of the western Pacific intraoceanic arcs.
They concluded that the crust production rates (between
50–95 km3 km−1 My−1) are somewhat higher than previous
estimates, but are still significantly less than those of
mid-ocean ridges.
For the approach outlined above, a crucial parameter is
the crustal thickness of an active intraoceanic arc. As this
parameter becomes better constrained, intraoceanic arc
crust rate estimates have historically increased with time
from the initial 30 km3 km−1 My−1 estimates of Reymer and
Schubert (1984) and Crisp (1984). Generally, it is thought
that the crust–mantle transition, and thereby the thickness of the crust, coincides with a significant change in
seismic-wave speeds, where primary-wave (P-wave) velocities increase from ~7 km s−1 to > 8 km s−1. In recent years,
it has become increasingly clear that arc thickness and
composition are hard to constrain remotely because the
petrologic crust–mantle transition does not correlate
with a significant increase in P-wave velocities (Jagoutz
and Behn 2013). This significantly hampers the calcula-
4 Long-term crust addition and magma production rates for arcs
and ridges are commonly expressed in units of volume (km3)
per km of arc length per unit of time, which is typically expressed
in 106 years or My (km3 km −1 My−1).
E LEMENTS
tion of crust production rates. Moreover, these estimates
yield only the time-averaged crust production rates, which
are often highly variable within the same arc system. For
example, in the Izu–Bonin–Mariana arc, the crust beneath
the northern Izu arc is three times as thick (~32 km) as
that beneath the central Bonin arc (~10 km) (Kodaira et
al. 2007). At a minimum, this corresponds to a factor of
three difference in magma production. More importantly,
quantifying the magma production rate (instead of crust
production rate) is even more complex due to the large
number of processes operating at subduction zones, which
can add or remove material from an arc and which need
to be taken into account (FIG. 1). Of these, there are three
main processes. First, rifting: the breaking up of a section
of arc crust, which is common in western Pacific intraoceanic arcs and can transport pieces of crust hundreds of
kilometers away from where they were emplaced. Second,
surficial and subduction erosion: the processes by which
arc crust is eroded or scraped away and transported into
the mantle on the downgoing plate (von Huene and Scholl
1991). Third, delamination: loss or sinking of dense lower
crust back into the mantle. The estimated bulk composition of the crust in intraoceanic arcs differs significantly
from that of mantle-derived melts as significant volumes of
material are lost due to the three processes described above
(Jagoutz and Schmidt 2012). Therefore crust production
rates significantly underestimate magma production rates.
In the following section, we briefly summarize existing
crust production rate estimates for active and accreted
intraoceanic arcs in the light of recent seismic, petrochemical, and geochronologic studies, and then attempt
to determine magma production rates for these arcs.
ACTIVE INTRAOCEANIC ARCS
Tonga–Kermadec Arc
The Tonga–Kermadec arc extends for 2500 km from
American Samoa to the North Island of New Zealand
(FIG. 2A). West-southwestward convergence of the Pacific
Plate beneath the Australian Plate is fast (20–25 cm y−1).
Subduction started around 50 Ma and was caused by a
subduction polarity-fl ip and the reactivation of a fossil
subduction zone (Whattam et al. 2008).
A detailed seismic tomographic image of the Tonga–
Kermadec subduction zone and a derived 2-D velocity
model (Contreras-Reyes et al. 2011) indicate that the arc’s
crustal thickness ranges between 7 and 20 km. The arc is
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A
B
C
D
A
B
The various
intraoceanic arcs
mentioned in the text:
(A) Tonga–Kermadec arc,
(B) Aleutian arc, (C) Izu–
Bonin–Mariana arc, and
(D) Lesser Antilles arc. Arrows
indicate direction of plate
motion. SOURCE : GOOGLE E ARTH
FIGURE 2
C
D
E LEMENTS
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inferred to be predominantly mafic in composition and
lacks the P-wave velocities of 6.0 ± 0.4 km s −1 that are
typical of continental crust. Crust production rates for the
Tonga–Kermadec arc range from 33 km3 km−1 My−1 (Reymer
and Schubert 1984) to 56 km3 km−1 My−1 (Dimalanta et al.
2002) (TABLE 1).
We recognize that there are other active intraoceanic arcs
(e.g. Kuriles, South Sandwich), but knowledge of these arcs
is generally poor and, thus, fruitful comparisons with the
aforementioned active arcs cannot be made.
Alaska–Aleutian Arc
There are several fossil intraoceanic arcs that have been
accreted to continents. These accreted arcs preserve a
nearly continuous rock record from the upper mantle to
the sediments overlying the volcanic sequence and serve
as a physical analog for understanding the structure and
processes that are inferred to have taken place in active
intraoceanic arcs.
The Alaska–Aleutian arc extends for ~3500 km from central
Alaska to the Kamchatka Peninsula along the northern
portion of the Pacific–North American plate boundary
(FIG. 2B). P-wave velocity models developed from wideangle seismic studies in the central Aleutian arc (e.g.
Holbrook et al. 1999), in combination with rare S-wave
arrivals, have been used to infer that the mid-to-lower crust
is composed of a mix of mafic and felsic rocks (Shillington
et al. 2013). The average crustal thickness is ~38.5 ± 2.9 km
(Janiszewski et al. 2013), which is unusually high for an
intraoceanic arc.
Based on an across-arc seismic reflection and refraction
survey, Holbrook et al. (1999) determined the volume of
material beneath the arc, assumed 55 Ma as the age of
arc formation, and calculated a magma production rate of
82 km3 km−1 My−1. Jicha et al. (2006) revised this production rate to 152–182 km3 km−1 My−1 based on a 46 Ma age
of arc inception and taking into account material lost over
the last 46 My via subduction and glacial erosion (TABLE 1).
Izu–Bonin–Mariana (IBM) Arc
The Izu–Bonin–Mariana (IBM) arc is located in the western
Pacific and extends from Japan to Guam (FIG. 2C). It has
been extensively studied over the last decade via dredging,
manned submersible sampling, and drilling because its
fore-arc preserves widespread exposures of lavas that were
generated during the initial stages of subduction. U–Pb
zircon and 40Ar/39Ar ages from fore-arc gabbros and basalts
spanning the entire 3000 km of the arc date back to 52 Ma
(Ishizuka et al. 2011). Despite the wealth of recent petrologic, seismic, and geochronologic data acquired for the
IBM arc, there have been no recent attempts to constrain
the rates of magma and crust production.
Lesser Antilles Arc
Subduction of the South American plate beneath the
eastern margin of the Caribbean plate has resulted in
the creation of the 850 km long Lesser Antilles volcanic
intraoceanic arc (FIG. 2D). Four features make the Lesser
Antilles arc distinctive. First, the arc bifurcates north of the
island of Martinique, with the active volcanoes forming the
inner arc to the west and inactive, limestone-capped islands
forming the outer arc to the east. Second, the convergence
rate (~2 cm y−1) is very slow (Wadge 1984). Third, unlike
the western Pacific arcs, which have very little sediment
accreted to them, the accretionary prism in the Lesser
Antilles arc is extremely thick (up to 20 km near Barbados).
Fourth, most of the arc may be underlain by a Cretaceous
island arc and/or a segment of the Caribbean plateau.
Wide-angle seismic data from a 280 km across-arc transect
indicates that the arc has a total average thickness of
~24 km with little variation along strike (Kopp et al. 2011).
There is no evidence for an ultramafic cumulate layer like
the one that is inferred to exist at the base of the IBM and
Aleutian arcs. The only estimate long-term magma production rate for the Lesser Antilles arc is that of Reymer and
Schubert (1984) (69 km3 km−1 My−1) (TABLE 1).
ACCRETED INTRAOCEANIC ARCS
Talkeetna Arc (Alaska, USA)
The Talkeetna arc was active during most of the Jurassic
Period (200–153 Ma) and was accreted to south-central
Alaska by the end of the Jurassic. In the Chugach Mountains,
a crustal cross section is exposed from the petrologic Moho
(crust–mantle boundary) through a sequence of lowercrustal gabbro norites to mid-crustal intermediate and felsic
plutonic bodies to a 7 km thick section of submarine and
subaerial volcanic and volcaniclastic deposits (Hacker et
al. 2008). Only 15–30% of the plutonic section survives
today. Even though the geochemical, petrologic, and
isotopic characteristics of the Talkeetna arc are now fairly
well constrained, the crust production rates have yet to
be established.
Kohistan Arc (Northern Pakistan)
The Kohistan arc in northern Pakistan is a Cretaceous
island arc that was wedged between the Indian and Asian
plates during collision around 50 Ma. Unlike the Talkeetna
arc, which is partially missing, the entire crustal section of
the Kohistan intraoceanic arc is preserved (e.g. Jagoutz and
Behn 2013), including a thick section of ultramafic rocks at
its base. Subduction-related magmatic activity in the mafic
lower crust and mid- to upper-crust intrusive and volcanic
rock spans 120 to 50 Ma (Bouilhol et al. 2013). The crust
production rate of the Kohistan arc is estimated to have
been ~48–72 km3 km−1 My−1 (Jagoutz and Schmidt 2012).
Jagoutz and Schmidt (2013), using an approach similar to
that outlined in the next section, estimated a significantly
higher magma production rate of 220–290 km3 km−1 My−1.
MAGMA PRODUCTION RATES:
WAYS FORWARD
As noted by Arculus (2004), the dynamic nature of a
subduction zone, in which the arc is losing material simultaneously with growth, makes it difficult to determine what
the magma production rate is or has been. This is an important point because most of the existing crust accumulation or magma production rate estimates for intraoceanic
arcs do not take into account the material that has been
removed from the arc since its inception. The total arc
magma production rate (Jarc) is the sum of the volume of
arc crust produced (V crust) and lost (V lost) divided by the
age of the arc (FIG. 1):
J arc = (V crust + V lost)/(age of arc)
Crust can be lost via any number of processes, including
rifting (V rift), subduction erosion (V sub ), delamination (V delam), and surficial erosion (V surf ). Low density
sediments, however, might be relaminated to the base of
the arc crust (V relam) and so counterbalance the surficial
and subduction erosion:
V lost = V rift + V sub + V delam + V surf – V relam
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TABLE 1
Age
(Ma)
Arc
Aleutians
Bulk
Compositiona
Existing
volume of
arc crust
(Vcrust )
(Jarc) Magma
production rate
(km3 km −1
My−1)
Source
33
[1]
[2]
55
4100
75
[3]
56
3750
67
[4]
55
3330
46
4100
0
1380
4100
0
0
BA (c)
61
4100
[5]
1500
89
152
[6]
1500
89
211
this study
45
1250
28
[1]
15
3000
200
[7] (i)
2950
60
[5]
BA (d)
4646
0
2080
4646
78
89
220
this study
45
1250
28
[1]
49
1480
30
[5]
BA (d)
1200
1875
2080
3075
78
23
160
this study
45
1750
39
[1]
45
2250
50
[5]
BA (d)
40
40
2839
2047
1040
4886
78
2750
BA (e)
60
3200
0
0
3200
78
B (f)
4000
100
BA
6000
209
this study
80
162
this study
[1]
33
1500
52
55
69
1950
27
Kohistan (g)
Crust production rate (km3
km −1 My−1)
34
52
Tonga–Kermadec
Surficial
erosion
(Vsurf )
2720
52
Lesser Antilles
Delamination
(Vdelam ) (i)
Subduction
erosion (Vsub ) (b)
2500
52
Mariana
Rifting
(Vrift )
75
49
Bonin
Volume lost
80
46
Izu
MAGMA PRODUCTION RATE (JARC) PARAMETERS FOR ACTIVE AND ACCRETED INTRAOCEANIC ARCS.
UNIT OF VOLUME IS KM3 PER KM OF ARC LENGTH.
[1]
56
0
2760
1320
78
[5]
77
157
this study
60
255
[8]
Notes: BA= basaltic andesite; B= basaltic; (a) Based on previous estimates or on updated geophysical parameters using the relationship
between vP and vP /vS and composition as outlined in Jagoutz and Behn (2013); (b) subduction erosion rates from Clift et al. (2009);
(c) based on an average crustal vP /vS 1.75 and a vP of 6.8; (d) based on Tatsumi et al. (2008); (e) based on Kopp et al. (2011); (f) based
on Contreras-Reyes et al. (2011); (g) duration of Kohistan arc magmatism is 100 Ma (150 Ma to 50 Ma); (h) delamination volumes are
calculated based on the estimated bulk composition (see Jagoutz and Schmidt 2013). BA: V arc = V delam ; B: V arc = 3V delam ; (i) magma
production rate calculated for the last 15 My of Izu–Bonin–Mariana arc history.
Sources: [1] Reymer and Schubert (1984); [2] Crisp (1984); [3] Holbrook et al. (1999); [4] Lizzaralde et al. (2002); [5] Dimalanta et al. (2002);
[6] Jicha et al. (2006); [7] Arculus (1999); [8] Jagoutz and Schmidt (2012, 2013)
We will now review each one of the processes contributing
to V lost.
Rifting
For intraoceanic arcs such as the Izu–Bonin–Mariana
system, which has experienced periodic rifting since the
Oligocene, the volume of the remnant arcs (V rift) must be
taken into account. For other arcs that are mostly intact,
such as the Aleutians, the V rift term is null.
Surficial Erosion, Subduction Erosion,
and Relamination
Von Huene and Scholl (1991) proposed that large volumes
of sediment and fore-arc crust are being subducted at many
of the Earth’s subduction zones, and they suggested that
arcs can be subdivided into two groups: accretionary
(V sub = 0) and erosive (V sub > 0). Erosive margins (e.g.
IBM, Tonga–Kermadec) are characterized by a fast subduction rate (>6.0 mm y−1), thin sedimentary trench fi ll
(<1.0 km thick), and abundant subduction erosion (V sub =
20–76 km3 km−1 My−1; Clift et al. 2009). Fore-arc crust may
E LEMENTS
also be eroded in accretionary subduction zones, but it is
difficult to quantify how much is being accreted versus
recycled into the mantle (von Huene and Scholl 1991).
In addition to accounting for the volume of crust removed
from the base of the arc via subduction erosion or delamination (see below), before surface erosion (V surf ) must also
be estimated. Unfortunately, the number of weathering
and erosion studies in intraoceanic arcs is limited. For the
Aleutian arc, Jicha et al. (2006) suggested that 6–9 km
of vertical thickness of the subaerial Finger Bay pluton
(2 × 2 km in outcrop) has been removed over the past
35 My, which corresponds to a maximum surface erosion
rate of ~1 km3 km−1 My−1. This erosion rate estimate was
extrapolated to the entire arc, though this may not be
representative of the surface erosion rate throughout the
Aleutians. However, most intraoceanic arcs are located in
areas that receive significant precipitation, and it has been
shown that these sites are, in turn, characterized by the
highest erosion rates on Earth (up to 1.5 km3 km−1 My−1)
(Gaillard et al. 2011).
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Additionally, it has been proposed that some of this
recycled material is less dense than the overlying mantle
wedge and, therefore, might rise as diapirs and be relaminated at the base of the arc (V relam ; Hacker et al. 2011).
Delamination
The observations and mass balance calculations of Jagoutz
and Schmidt (2013) indicate that the chemical and petrologic composition of the middle to lower crust in intraoceanic arcs is often not in equilibrium with the upper
mantle. They suggest that foundering or delamination of
dense, unstable mafic and ultramafic cumulates near the
Moho is a volumetrically important and underappreciated process. They also suggest that the principal formation mechanism for evolved melts in intraoceanic arcs is
fractionation and that the volume lost to delamination
(V delam) is approximately twice that of the remaining arc
crust (V crust). In the Kohistan arc, the combination of
fractionation of primitive, high-Mg basalt and delamination of ultramafic cumulates reproduces the intermediatecomposition plutons found in the middle crust (Jagoutz
and Schmidt 2013).
Unfortunately, many of these processes are poorly
constrained or entirely unconstrained. Nevertheless, we
can exploit the fact that some of the processes resulting in
loss of material also produce a net change in composition
of the bulk arc crust. A simple mass balance determines
the concentration of an element i in the bulk arc crust:
The lowest magma production rate is for the Tonga–
Kermadec arc (157 km3 km−1 My−1) and the highest is
for the Izu arc (220 km3 km−1 My−1). Intermediate values
come out for the Aleutians (211 km3 km−1 My−1), Bonin
(160 km 3 km−1 My−1), Mariana (209 km 3 km−1 My−1),
and Lesser Antilles (162 km3 km−1 My−1) (FIG. 3; TABLE 1).
Similarly, Jagoutz and Schmidt (2013) calculate V lost and
V delam and suggest that the rate of magma production in
the Kohistan arc was ~220–290 km3 km−1 My−1.
SIGNIFICANCE OF REVISED MAGMA
PRODUCTION RATES
Our results indicate that magma production in intraoceanic arcs has been profoundly underestimated because
reworking processes have not been properly accounted for.
The crust production rate for the Sierra Nevada during
the fl are-up periods is not unusual and is comparable
to crust production rates observed in intraoceanic arcs
(FIG. 3). The unusual stages are the intermittent quiescent
periods. Similar quiescent periods are currently observed
in the Andean margin between the northern, central, and
southern volcanic zones where magmatism has essentially
ceased due to flat-slab subduction (Stern 2004).
Our reanalysis further suggests that despite the difference
in tectonic setting, the rates at which magma is produced at
intraoceanic arcs (157–290 km3 km−1 My−1) and mid-ocean
ridges (~280 km3 km−1 My−1) is similar (FIG. 3). The contri-
(x) Ciarc = Cimantle melt – (1 – x) Cilost
where Ciarc, Cimantle melt, and Cilost represent the concentration of an element of interest (e.g. SiO2 ) in the bulk arc
crust, the mantle derived melts and in the material lost,
respectively, and x indicates the relative mass proportions.
To outline the importance of these different mechanisms
for arc magma production rates, we used seismic velocity
measurements to roughly constrain the bulk compositions of different segments of the arc crust. Specifically,
we employed the relationship between vP /vS and vP (where
vP = P-wave velocity; vS = S-wave velocity) and chemical
composition, as calculated by Jagoutz and Behn (2013).
Given the uncertainty in relating seismic velocities to
crustal composition, we roughly divided arc crust into
two compositional groups based on the seismic properties: basaltic (e.g. SiO2 < 52 wt%) and basaltic–andesitic
(SiO2 > 52 wt%) (TABLE 1). Based on the results of Jagoutz
and Schmidt (2013), we estimated that the V delam is about
equal to the total crust produced (V crust + V rift + V surf +
V sub) for basaltic–andesitic bulk composition, whereas in
basaltic arcs the V delam is 1/3 of the total crust produced (see
footnote, TABLE 1). We also made the assumption that the
material removed by rifting and by surficial and subduction
erosion is comparable in composition to that of the bulk arc
crust. Due to the complete lack of constraints on V relam, we
simply ignored this process. We also decided to ignore the
potential volume of preexisting oceanic crust because in
exposed arc terranes (e.g. Talkeetna and Kohistan arcs) as
well as in the IBM fore-arc (Reagan et al. 2010), old oceanic
basement has not been observed. This might relate either
to how subduction initiates and new suprasubduction zone
crust is formed or to the preexisting oceanic crust being
completely assimilated and intruded so that it is unrecognizable. Clearly, a better understanding the mechanisms of
subduction initiation and the fate of a possible preexisting
oceanic crust in arcs might alter the results presented here.
Using the parameters outlined above, the total arc
magma production rates for intraoceanic arcs can be
constrained better. Our magma production rate estimates
for active intraoceanic arcs vary significantly (FIG. 3).
E LEMENTS
Scatter plot of normalized volume versus age of each
arc. Volumes for various arcs are normalized per
kilometer of arc length. Yellow and red symbols are average crust
production estimates that neglect potential reworking processes;
blue symbols are magma production estimates (data sources given
in Table 1). The 255 km3 km −1 My−1 rate is for the Kohistan arc
(Jagoutz and Schmitz 2013). The green field indicates magma
production (i.e. crust production) along the global oceanic
spreading centers (data after Bird 2003), assuming 5–7 km thick
oceanic crust. Black lines are contours of constant production.
The orange fields are the flare-up and quiescent periods of the
Cretaceous Sierra Nevada arc (Ducea et al. 2015). The flare-up
periods have crust production rates comparable to those observed
in active intraoceanic arcs. The quiescent periods, however, have
extremely low production rates and indicate termination of
magmatic activity. The difference between magma and crustal
production estimates illustrates the importance of reworking
mechanisms along suprasubduction zones.
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FIGURE 3
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bution from intraoceanic arcs to the global magma generation budget and Earth’s loss of internal heat has been
underappreciated and needs to be reevaluated. As the total
length of subduction zones on Earth (51,310 km) is about
75% of the total length of ridges on Earth (67,338 km;
Bird 2003), our revised magma production rates of
~157–290 km3 km−1 My−1 indicate that the contribution
of heat loss from subduction zone magmatism should be
on the order of 1.2–3.3 terawatts. Accordingly, the heat
loss due to magmatism along ridges and subduction zones
on Earth is, collectively, ~5.2–7.3 terawatts (~13–17% of
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