Terrestrial heat flow in the Malawi Rifted Zone, East Africa: Implications for tectono-thermal inheritance in continental rift basins

Journal of Volcanology and Geothermal Research 387 (2019) 106656

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Terrestrial heat flow in the Malawi Rifted Zone, East Africa: Implications for tectono-thermal inheritance in continental rift basins Emmanuel A. Njinju a,⁎, Folarin Kolawole b, Estella A. Atekwana c, D. Sarah Stamps a, Eliot A. Atekwana c, Mohamed G. Abdelsalam d, Kevin L. Mickus e a

Department of Geosciences, Virginia Tech, Blacksburg, VA, USA School of Geology and Geophysics, University of Oklahoma, Norman, OK, USA Department of Geological Sciences, University of Delaware, Newark, DE, USA d Boone Pickens School of Geology, Oklahoma State University, Stillwater, OK 73019, USA e Department of Geography, Geology and Planning, Missouri State University, Springfield, MO, USA b c

a r t i c l e

i n f o

Article history: Received 25 March 2019 Received in revised form 29 July 2019 Accepted 31 July 2019 Available online 07 August 2019 Keywords: Tectono-thermal inheritance Aeromagnetic data Curie point depth Heat flow Malawi Rifted Zone Geothermal exploration

a b s t r a c t Rift basins show the widest range of heat flow values compared to passive continental margins and intracratonic sag basins. Elevated heat flow is often used to infer asthenospheric upwelling or subsurface magmatic activity for magma-assisted rifts. Despite the lack of surface volcanism for magma-poor rifts, it is possible that magmatic bodies will be present at lower and mid crustal levels but yet to breach the surface, thus necessitating consideration of heat flow from tectono-thermal inheritance in addition to more recent magmatism. We use aeromagnetic data to investigate the terrestrial heat flow distribution in the Malawi Rifted Zone (MRZ) in East Africa defined as the Neogene Malawi Rift and surrounding Permo-Triassic Karoo rift basins. We use the twodimensional power-density spectrum technique to estimate the Curie point depth (CPD), geothermal gradient, and heat flow beneath the MRZ. Our results reveal predominance of shallow CPDs (18–20 km), high geothermal gradients (29–32 °C/km) and elevated heat flow (70–82 mW m−2) within the Karoo rift basins. Along the Malawi Rift, geothermal gradient (25–27 °C/km) and heat flow (60–66 mW m−2) are generally lower except pronounced high heat flow (70–82 mW m−2) at the Rungwe Volcanic Province (RVP) in the north, and at the central riftsegment where the rift overprints Karoo rift basins. The surrounding cratonic blocks show deeper CPDs (24–27 km), lower geothermal gradients (22–24 °C/km) and lower heat flow (53–63 mW m−2). We interpret that the elevated heat flow in the Karoo rift basins are related to residual heat flow from the Permo-Triassic rifting-phase, replenished by later Jurassic-Cretaceous diking events. Apart from the thermal anomaly beneath the RVP, which is due to Neogene magmatism, other important targets for geothermal explorations in the MRZ are the residual thermal anomalies in the Karoo rift basins. We infer that in areas of active magma-poor rifting, pronounced heat flow may not only occur in regions of asthenospheric upwelling, but may concentrate in the ancient magmatic rift segments. Our findings have implications for crustal stretching, tectono-thermal inheritance, and geothermal energy potentials in East Africa and other segments of magma-poor rifting. © 2019 Elsevier B.V. All rights reserved.

1. Introduction Terrestrial heat flow in any continental or oceanic province is strongly related to the age of the province (e.g., Lee and Uyeda, 1965; Pollack, 1982; Polyak and Smirnov, 1968; Sclater and Francheteau, 1970) and pronounced residual heat flow may be preserved for millions of years after the cessation of deformational activity (e.g., Hu et al., 2000). For continental regions, the last tectono-thermal event, distribution of heat-producing elements, and erosion are the main factors that determine the surface heat flow (Pollack, 1982; Sclater et al., 1980). ⁎ Corresponding author. E-mail address: [email protected] (E.A. Njinju).

https://doi.org/10.1016/j.jvolgeores.2019.07.023 0377-0273/© 2019 Elsevier B.V. All rights reserved.

Consequently, cratonic blocks have significantly lower heat flow compared to the surrounding Proterozoic and Early Paleozoic orogenic belts along which continental rifts subsequently develop (Kearey et al., 2013; Lysak, 1987; Nyblade, 1997; Nyblade et al., 1990; Sclater et al., 1980). Compared to passive continental margins and intracratonic sag basins, rift basins show the greatest heat flow variability (widest range of heat flow values; Figs. 1 a–d). Heat flow in continental rifts varies with the rift type (magmatic or magma-poor), rifting stage, intra-rift structures and the age of recent tectono-thermal activity in the rift zone (Lysak, 1987). Regional thermal anomalies are mainly associated with conductive heat transfer from the asthenosphere, whereas, local thermal anomalies are often associated with magma bodies, radioactive

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Fig. 1. (a) Global distribution of continental basins (source: Tellus™ Sedimentary Basins of the World Map) overlaid with heat flow measurement locations (International Heat Flow Commission, Southern Methodist University National Geothermal Data System). Histogram of heat flow distribution for (b) Rift basins, (c) Passive Margins, and (d) Intracratonic Sag basins. The histograms (normalized scales for the three basin types) show that among these basins, rifts show the greatest heat flow variability (long tail of the histogram plot compared to passive margins and intracratonic sag basins). The blue box labeled 2A shows the location of Fig. 2a. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

elements in the upper crust, and by hydrothermal fluid convection along permeable multiscale fracture/fault systems (e.g., Lysak, 1987). However, temporal dissipation of heat flow on earth follows an exponential distribution (e.g., Pollack, 1982; Sclater et al., 1980). In continental tectonic provinces, thermal decay of the heat flow takes 800 million years to reach a constant value of 46 mW m−2 (Sclater et al., 1980), such that the mean heat flow through provinces of Mid-Permian age is about 58 ± 12 mW m−2, and at least 73 ± 10 mW m−2 through provinces younger than Late Eocene (Chapman and Pollack, 1975; Polyak and Smirnov, 1968). However, there are certain geologic and tectonic settings where the regional heat flow may depart significantly from the appropriate global age-group mean (Pollack, 1982). Given the wide range of heat flow values in rift basins (Fig. 1b), there is the need to consider the role of tectono-thermal inheritance in addition to more recent magmatism in heat flow variation in rift basins as this may significantly perturb the normal thermal subsidence of a cooling rifted continent. Lithospheric thermal perturbation is fundamental to the process of rifting in both oceanic and continental settings at various stages of the rift development (e.g., Chapman and Pollack, 1977; Leroy et al., 2010; Leseane et al., 2015; Lubimova and Polyak, 1969; Lysak, 1987; Morgan, 1982, 1983; Nyblade et al., 1990; Rooney et al., 2012; Wheildon et al., 1994). Along rifted margins, the presence of high thermal gradients can result in varying along strike rift structure and deformational styles due to thermal weakening of the lithosphere and triggering of crustal buoyancy forces (Bellahsen et al., 2013; Leroy et al., 2010). Also, anomalous heating of the lithosphere and the associated crustal uplift can be associated with rift-related asthenospheric upwelling which can induce “rift-push” forces to drive continental extension (Mareschal, 1983; Mareschal and Gliko, 1991; Rooney et al., 2012). Furthermore, higher heat flow coming from subsurface magmatic regions formed at the onset of rifting can provide the initial mechanical weakening of the lithosphere to facilitate the progression of continental rifting (Bialas et al., 2010; Buck, 2004). Aside from the influence of thermal anomalies at the later stages of continental rifting (Bastow et al., 2005; Bellahsen et al., 2013; Lysak, 1987), thermally assisted amagmatic crustal deformation appears to play a fundamental role in the process of rift nucleation right at the onset of continental rifting (Leseane et al., 2015). It has also been demonstrated that high surface heat flow along amagmatic continental rifts may not develop until after crustal thinning has ceased (Morgan, 1983), but, residual heat flow along inactive continental rifts can be preserved for N60 Ma after cessation of deformational activity (Hu et al., 2000). In addition

to heat flow due to plate boundary tectonic activity, thermal anomalies have been correlated with seismicity and fault reactivation in stable intraplate settings (Holford et al., 2011). Although information about the distribution of thermal anomalies in rifted zones can provide insight into the processes of continental rifting, such data are also highly useful for geothermal energy development. The geothermal energy potential of several countries (e.g., Ethiopia, Kenya, Uganda) traversed by the East African Rift System (EARS) has been assessed (Kiplagat et al., 2011; Teklemariam, 2008), and progress has been made in the development of the resource (Bertani, 2005; Fridleifsson, 2001; Kiplagat et al., 2011). Along the EARS and other geothermally active regions worldwide, terrestrial heat flow variations have been assessed using measurements of surface hot spring temperature (Ferguson and Grasby, 2011; Lonsdale and Becker, 1985), and in mining, hydrocarbon and groundwater bore holes (Chapman and Pollack, 1977; Hu et al., 2000; Nyblade et al., 1990; Wheildon et al., 1994) and lake bed sediments (Hart and Steinhart, 1965; Lubimova and Polyak, 1969; von Herzen and Vacquier, 1967). These techniques can be expensive or may produce limited coverage due to inaccessibility of sampling locations. However, relatively less expensive techniques, such as Curie point depth (CPD) estimations from aeromagnetic data, have produced reliable results of the regional terrestrial heat flow distribution in numerous regions (Aboud et al., 2011; Afshar et al., 2017; Arnaiz-Rodríguez and Orihuela, 2013; Leseane et al., 2015; Nwankwo and Shehu, 2015). In this contribution, using aeromagnetic data, we present for the first time, the detailed crustal thermal structure of the entire Malawi Rifted Zone (MRZ), consisting of the Neogene Malawi Rift, and the PermoTriassic (Karoo) Rukwa Rift, Luangwa Rift, Shire Graben, Ruhuhu Trough, Maniamba Trough and the Zambezi Rift (Figs. 2a–c). Our results reveal a striking predominance of elevated heat flow within the PermoTriassic Karoo rift basins in the MRZ. The currently active Malawi Rift does not show any systematic increase in heat flow beneath the rift axis except for the Rungwe Volcanic Province (RVP) in the north, and the central segment where the rift overprints the Karoo Maniamba Trough. Our findings reveal that in areas of active magma-poor continental rifting, elevated heat flow will not only concentrate at the isolated active magmatic zones, but may also focus mostly in ancient magma-assisted and/or thinned-crust rift segments. Our findings provide insights into rift-related heat flow distribution, the implications for crustal stretching, tectono-thermal inheritance, and geothermal energy potentials in areas of magma-poor rifting.

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Fig. 2. (a) Digital Elevation Model (DEM) extracted from the Global 30 Arc Second Elevation Data (GTOPO30) showing the southern part of the East African Rift System (EARS) and the Eastern and Western branches of the Cenozoic EARS (in red, 1) and the Mesozoic Karoo rifts (in yellow, 2). TR = Tanganyika Rift. RR = Rukwa Rift. MR = Malawi Rift. LR = Luangwa Rift. UG = Urema Graben. UR = Usangu Rift. ZR = Zambezi Rift. RT = Ruhuhu Trough. MT = Maniamba Trough. The white box indicates the Malawi Rifted Zone (MRZ). (b) Shuttle Radar Topography Mission (SRTM) Digital Elevation Model (DEM) of the MRZ showing the border faults and the surrounding Paleozoic-Mesozoic Karoo rift basins. Blue contour lines show water depth within Lake Malawi. RVP = Rungwe Volcanic Province. (c) Tectonic map showing the exposures of the Precambrian and Paleozoic-Mesozoic units around the MRZ. Modified after Fritz et al. (2013) and Laó-Dávila et al. (2015). T = Txitonga Group. PM = Ponta Messuli Complex. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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2. Geological setting The MRZ is located in the southern segment of the EARS (Fig. 2a), a N600,000 km2 region which includes the Neogene Malawi Rift and Rukwa Rifts, and the nearby Permo-Triassic rift basins such as the Luangwa Rift, Shire Graben, Zambezi Rift, Ruhuhu Trough and Maniamba Trough (Fig. 2b). The entire region is underlain by Proterozoic orogenic belts (e.g., Ubendian, Usagaran and Mozambique Belts) that surround the larger Tanzanian Craton and the smaller Bangweulu cratonic block and Niassa Craton (Fig. 2c; Daly, 1988; Fritz et al., 2013; Lenoir et al., 1994; Sarafian et al., 2018). The Proterozoic orogenic belts consist of various basement terranes that are separated by lithospheric-scale shear zones of distinct structural and petrological architecture (Boven et al., 1999; Daly, 1988; Heilman et al., 2019; Kolawole et al., 2017; Sarafian et al., 2018). Studies have highlighted the presence of relics of Proterozoic subduction zones along some of the orogenic belts (Boniface and Schenk, 2012; Boniface et al., 2012, 2014; Ring, 1993, 1994; Ring et al., 2002). 2.1. Permo-Triassic rifting In the Phanerozoic, the region of the MRZ first underwent an episode of continental extension during the Permo-Triassic breakup of Pangaea, known as the “Karoo” rifting episode, during which fault-bounded graben and half graben rift basins developed along the Proterozoic orogenic belts (Castaing, 1991; Daly et al., 1989; Delvaux and Hanon, 1991). During this rifting event, thick packages of fluvio-lacustrine sediments and coal units, which commonly thicken towards the rift border faults, were deposited unconformably on the crystalline basement (Castaing, 1991; Morley et al., 1999). The Karoo intracontinental rift basins generally trend NNE-SSW, NW-SE and E-W (Castaing, 1991; Delvaux et al., 1992; Morley et al., 1992) and the development of the basin-bounding faults were controlled by the fabrics of the Precambrian basement terranes and the associated shear zones (Castaing, 1991; Daly et al., 1989). In the MRZ, the Karoo rift basins include the Luangwa Rift, the Ruhuhu and Maniamba (Metangula Basin) Troughs, the Shire Graben and the Zambezi Rift (Figs. 2a–b; Castaing, 1991; Morley et al., 1999). The later stages of the Karoo rifting recorded widespread magmatic activity, which included the emplacement of thick volcanic sequences and the intrusion of dikes and sills in both the basement and sedimentary units (Castaing, 1991). The Karoo rifting event was concluded by an episode of basin inversion during which the Karoo sediments were folded (Castaing, 1991). 2.2. Middle Jurassic to early Cretaceous rifting After the Karoo rifting event, the MRZ recorded another transient phase of extension in the Middle Jurassic to Early Cretaceous, during which alkaline igneous intrusions (ring complexes) were emplaced in the Shire graben area of the southern Malawi Rift, which is known as the Chilwa Alkaline Province (Castaing, 1991). This Cretaceous extensional event was associated with the deposition of fossiliferous red, fluviatile sandstone, and siltstones referred to as the “Dinosaur Beds” within the Karoo basins (Jacobs et al., 1993; Ring, 1994). 2.3. Cenozoic rifting In the Cenozoic Era, the MRZ began to undergo yet another phase of extension that affected most of eastern Africa. This resulted in the reactivation of the border faults of some of the Karoo rift basins (e.g., the Rukwa Rift, North Malawi Rift). Subsidence along the faults allowed the resumption of sediment deposition within the basins (Castaing, 1991; Morley et al., 1999). The Rukwa and North Malawi Rift propagated towards each other and currently exhibit a coupled structure that is modulated by the Precambrian basement fabrics (Heilman et al., 2019). In these rift segments seismogenic fault ruptures

concentrate at the rift transfer and half-graben hinge zones (e.g., Camelbeeck and Iranga, 1996; Delvaux and Hanon, 1991; Heilman et al., 2019; Kolawole et al., 2018). The Malawi Rift propagated southwards (Calais et al., 2006; Laó-Dávila et al., 2015; Saria et al., 2014; Specht and Rosendahl, 1989; Stamps et al., 2008) across the Ruhuhu and Maniamba Karoo basins and terminated at the Shire Graben in the south. Also, the extension resumed in the Shire Graben with the development of a new border fault and accumulation of alluvial sediments (Castaing, 1991). Although the Luangwa Rift, Ruhuhu Trough, Maniamba Trough and the Zambezi Rift have hosted recent earthquakes (Nyblade and Langston, 1995), these basins appear to be relatively inactive and may not have accommodated significant strain driving the ongoing Neogene extension. The Usangu rift basin is weakly extended, and it is not clear if the rift developed in the Permo-Triassic (Mbede, 2002) or in the Neogene period (Harper et al., 1999). The Neogene episode of rifting that formed the present-day Malawi Rift was also accompanied by magmatic activity limited to the RVP (Fig. 2b). The RVP is located within the accommodation zone between the Rukwa Rift and North Malawi Rift (Ebinger et al., 1989; Fontijn et al., 2012). The RVP has recorded several episodes of volcanic activity since the initiation of Neogene extension (Fontijn et al., 2010, 2012) and the recent clustering of seismic activity in the area may suggest active tectonic-related processes beneath the volcanic province (Accardo et al., 2017; Camelbeeck and Iranga, 1996; Delvaux and Hanon, 1991). 3. Data and methods 3.1. Aeromagnetic data The aeromagnetic data used in this study is a merged grid obtained from the Council of Geosciences, South Africa and includes aeromagnetic data collected by the Government of Malawi between 1984 and 1985. The data were acquired along E-W trending flight lines with a 1 km line spacing, 10 km tie lines, and 120 m terrain clearance. The International Geomagnetic Reference Field (IGRF) model, which represents the geomagnetic field due to hydrodynamic processes in the outer core was removed from the aeromagnetic data in order to obtain the total magnetic intensity (TMI) data due to crustal magnetization. No data were collected over Lake Malawi. Satellite TMI data from the EMAG2 (Earth Magnetic Anomaly Grid) dataset (NOAA website doi: https://doi.org/10.7289/V5H70CVX) were used to supplement coverage over the lake. The EMAG2 data are publicly available airborne, shipboard, and satellite data merged into a 2 arcsecond grid that is 4 km upward continued above the Earth's surface. These data were downward continued to 120 m and merged with the TMI data at a resolution of 0.03° x 0.03° (Fig. 3). 3.2. Curie point depth analysis The aeromagnetic data are used to estimate the depth to the bottom of magnetic sources, which could include the CPD beneath the MRZ. These CPDs are subsequently used as a proxy for the determination of the geothermal gradients and heat flow of the MRZ. The CPD is the depth at which magnetic minerals within the crust reach their Curie temperature, or temperature at which magnetic minerals lose their ferromagnetization or ability to become permanently magnetized. The Curie temperature is generally assumed as 580 °C, corresponding to the Curie temperature of pure magnetite (Hunt et al., 1995; Ross et al., 2006). However, there are wide varieties of magnetic minerals within the crust with oxides being the most common. These oxides range from pure iron oxide (magnetite) to different iron oxide minerals that contain copper, cobalt, manganese, magnesium and/or nickel (Hunt et al., 1995). The Curie temperature for these oxides range from 440 to 580 °C, but magnetite is the most common mineral in the continental crust (Hunt et al., 1995). However, one cannot rule out that bottom of the magnetic sources are simply the bottom of a magnetic body that is

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Fig. 3. Map of total magnetic intensity of the aeromagnetic data used for estimating the Curie Point Depth (CPD) beneath the Malawi Rifted Zone using two-dimensional (2D) radially averaged power spectrum analysis. Boxes show an example of 50 % overlapping of the 110 × 110 km (1° × 1°) windows during the calculation of the power spectrums. Black box represents the location of the 110 × 110 km window used for the generation of the power spectrum curves in Fig. 4. The longer black dashed lines represent the outline of the Malawi Rift.

above the Curie isothermal point. Additional geological and geophysical data are needed to aid in determining if the bottom depths truly represent the CPD. There are a variety of approaches to estimate the depth to the base of magnetic sources. They can be broadly divided into two techniques: 1) power spectrum methods and 2) forward and inversion methods. The power spectrum methods include: (1) the spectral peak method, which determines the position of the observed spectral peak to be a function of depth to the top (Zt) and base (Zb) of the magnetized layer (Connard et al., 1983; Ravat et al., 2007; Ross et al., 2006; Shuey et al., 1977; Spector and Grant, 1970); (2) centroid method which determines the relationship between the slopes of plots of the power spectrum of the magnetic anomalies against wavenumbers to represent the depths to the base of the magnetized crust (Bhattacharyya and Leu, 1975, 1977; Okubo et al., 1985; Tanaka et al., 1999); and (3) the modifiedcentroid fractal method which is based on scaling distribution of the magnetic sources (Bansal et al., 2011, 2013; Maus and Dimri, 1996). Forward modeling or inversion methods can be used to model isolated magnetic anomalies for two-dimensional bodies (Byerly and Stolt, 1977;

Hong, 1982; Mickus, 1989) or a three-dimensional region (Hussein et al., 2013) to determine the depth of the magnetic sources. We used the centroid method because it gives reliable estimates of the CPD with less depth errors compared with the other methods (Ravat et al., 2007). In this study we determined the CPD using the power spectrum based centroid method, which is based on the examination of the shape of isolated magnetic anomalies (Bhattacharyya and Leu, 1975, 1977) and the study of the statistical properties of the magnetic anomalies (Spector and Grant, 1970). The centroid method is applied in the frequency domain and relates the 2D radially averaged power spectrum of magnetic anomalies to the depth to the base of the magnetized sources (Hussein et al., 2013; Manea and Manea, 2011; Okubo et al., 1985; Shuey et al., 1977; Spector and Grant, 1970; Tanaka et al., 1999; Tanaka and Ishikawa, 2005). The 2D radially averaged power spectrum of magnetic anomalies are calculated in ~110 × 110 km (1° × 1°) windows that are overlapped by 50% on all sides in order to increase the resolution of our results (Tanaka et al., 1999), and also to reduce edge effects due to the Gibbs phenomenon that occurs when applying the Fast Fourier transform. Utilization of a smaller window size is a

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fundamental error in the application of the spectral methods for aeromagnetic interpretation (Ravat et al., 2007) as smaller windows may have overlapped deeper and shallower spectral components on the spectral curves and are therefore incapable of distinctly determining the CPD. If the magnetized sources have CPDs greater than L/2π, where L is the length of the diagonal of a data window (e.g., L = 155.5 km for a 110 × 110 km window), they may not be properly resolved by the spectral method (Shuey et al., 1977). Therefore, a data window of 110 × 110 km (1° × 1°) will satisfactorily resolve CPD up to depths of 23.6 km. However, during our analysis, several windows with CPDs N23.6 km are re-computed with an increased data window size of ~220 × 220 km (2° × 2°) to insure that these greater depths are true. The centroid method to determine the CPD, or depth to the base of the magnetized source, consists of first calculating the 2D radially averaged power spectrum of magnetic anomalies in the frequency domain. Then, the spectrum is plotted against wavenumber and the depth can be determined in two steps: (1) estimate the depth to the top (Zt) of magnetized body and (2) estimate the depth to the centroid (Zc) of the deepest magnetized source (Okubo et al., 1985; Tanaka et al., 1999). In this method, the depth to the top of the magnetized source (Zt) is related to the power spectral density (P(k)) as follows (Spector and Grant, 1970):

The centroid method has several limitations, including: (1) the calculated CPD values is representative of the entire 1° × 1° sub-region, as the depths may be an average of shallow and deep CPD values within the sub-region (Hussein et al., 2013); and (2) uncertainties about the nature of magnetization at depth. It is possible that the calculated CPD values may represent the depth to lithological contact and not necessarily the depth at which magnetization is lost at high temperatures (Ross et al., 2006). In spite of these limitations, the centroid method has been successfully used to determine the CPD in different regions such as: eastern and southern Asia (Tanaka et al., 1999), central-southern Europe (Chiozzi et al., 2005), the Japenese Islands (Tanaka and Ishikawa, 2005), Mexico (Manea and Manea, 2011), the Death Valley, California (Hussein et al., 2013), and the Sokoto basin, Nigeria (Nwankwo and Shehu, 2015). To determine the accuracy of the calculated CPD values, the statistical error is calculated for each window from the ratio of the standard deviation of the slopes of three linear fits, to the range of the wavenumber used when determining the gradients for Zt and Zc from the spectral curves (Chiozzi et al., 2005). The error ranges in this study are found to be from 0.02 to 0.15 km for Zt, and from 1.8 to 3.1 km for Zc.

  1=2 In PðkÞ ¼ A− j k j Zt

The CPDs calculated from the aeromagnetic data analysis can be used as a proxy for calculating the geothermal gradient (dT/dz) and heat flow (q) using the following relationships (Tanaka et al., 1999; Ross et al., 2006):

ð1Þ

where P(k) is the radially averaged power spectrum, k is the wavenumber (2π rad/km), A is a constant and Zt is the depth to the top of the magnetized source. Thus, Zt is determined by plotting In(P(k)1/2) against the wavenumber (k) (Fig. 4a) and calculating the slope of the linear fit at the higher wavenumber portions (e.g., 0.7–1.4 rad/km) of the spectral curve. Tanaka et al. (1999) advocates fitting the slope to the higher wavenumber portions of the spectral curve arguing that the linearized equation for the depth to the top is valid for wavelengths greater than the thickness of the layer. Similarly, the depth to the centroid of the magnetized source (Zc) is related to the power spectral density P(k) as follows (Okubo et al., 1985):   1=2 In PðkÞ =k ¼ B− j k j Zc

ð2Þ

where B is a constant and Zc is the depth to the centroid of the magnetized source. Thus Zc is determined by plotting In(P(k)1/2/k) against the wavenumber (k) (Fig. 4b) and calculating the slope of a linear fit at the lower wavenumber portions (e.g., 0.10–0.31 rad/km) of the curve. The CPD or depth to the base of the magnetized crust, Zb is given by (Okubo et al., 1985): Zb ¼ 2Zc −Zt

ð3Þ

3.3. Geothermal gradient and heat flow estimation from Curie point depths

dT=dz ¼ θc=Zb

ð4Þ

where θc is the Curie temperature. We use the Curie temperature of magnetite (580 °C), the most abundant magnetic mineral in the crust (Hunt et al., 1995; Ross et al., 2006). The temperature of 580 °C represents the highest Curie temperature of the commonly found magnetic minerals in the crust. The Curie temperature depends on the magnetic mineralogy in the crust where the most common magnetic minerals (Gilder and Le Goff, 2008) in the crust are titanomagnetites (Fe3 −xTixO4), which is a solid solution series that has end members magnetite (Fe3O4) and ulvospinel (Fe2Ti1O4). Increased titanium content in titanomagnetite systematically lowers the spontaneous magnetization and the Curie temperature (Hunt et al., 1995). Thus any heat flow calculations using the Curie temperature of magnetite will determine the highest heat flow values in a region. The determination of the geothermal gradient is based on the following assumptions: (1) there are no heat sources or heat sinks between the Earth's surface and the CPD; (2) the surface temperature is 0 °C, and (3) the geothermal gradient (dT/dz) is constant. Based on these assumptions, a first-order estimate of the regional heat flow can be determined, using the one-dimensional (1-D) conductive heat flow equation. The 1-D Fourier's Law of conductive heat flow equation that

Fig. 4. Example of the two-dimensional (2D) radially averaged power spectrum analysis curve (a) for the depth to the top (Zt) of the magnetic source and (b) for the depth to the centroid (Zc) of the magnetic source that were used to estimate the Curie Point Depth (CPD) for the 110 × 110 km (1° × 1°) black box shown in Fig. 3.

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Table 1 Estimated geothermal parameters. Latitude (degrees)

Longitude (degrees)

Curie point depths (km)

Geothermal gradient (°C/km)

Heat flow (mW m−2)

−8.5 −8.5 −8.5 −8.5 −8.5 −8.5 −8.5 −8.5 −8.5 −8.5 −8.5 −9.5 −9.5 −9.5 −9.5 −9.5 −9.5 −9.5 −9.5 −9.5 −9.5 −9.5 −10.5 −10.5 −10.5 −10.5 −10.5 −10.5 −10.5 −10.5 −10.5 −10.5 −10.5 −11.5 −11.5 −11.5 −11.5 −11.5 −11.5 −11.5 −11.5 −11.5 −11.5 −11.5 −12.5 −12.5 −12.5 −12.5 −12.5 −12.5 −12.5 −12.5 −12.5 −12.5 −12.5 −13.5 −13.5 −13.5 −13.5 −13.5 −13.5 −13.5 −13.5 −13.5 −13.5 −13.5 −14.5 −14.5 −13.5 −14.5 −14.5 −14.5 −14.5 −14.5

31.5 32.0 32.5 33.0 33.5 34.0 34.5 35.0 35.5 36.0 36.5 31.5 32.0 32.5 33.0 33.5 34.0 34.5 35.0 35.5 36.0 36.5 31.5 32.0 32.5 33.0 33.5 34.0 34.5 35.0 35.5 36.0 36.5 31.5 32.0 32.5 33.0 33.5 34.0 34.5 35.0 35.5 36.0 36.5 31.5 32.0 32.5 33.0 33.5 34.0 34.5 35.0 35.5 36.0 36.5 31.5 32.0 32.5 33.0 33.5 34.0 34.5 35.0 35.5 36.0 36.5 31.5 32.0 36.5 31.5 32.0 32.5 33.0 33.5

24.0 23.2 21.8 22.8 24.7 23.7 25.5 23.0 17.8 19.4 19.2 26.0 27.3 22.9 23.2 24.3 20.9 24.3 20.0 20.7 19.5 25.2 24.3 27.0 25.0 26.6 26.7 26.7 20.7 18.1 18.0 18.1 20.9 23.5 22.7 24.1 25.4 22.9 24.0 22.6 19.7 21.9 20.9 22.1 24.1 21.3 23.2 23.3 22.8 18.8 22.7 19.6 22.7 23.4 21.8 24.1 23.0 24.1 23.7 24.5 24.6 23.2 26.0 24.1 22.2 22.5 24.5 22.6 22.5 24.5 22.6 24.7 24.5 22.2

24.2 25.0 26.6 25.4 23.5 24.5 22.7 25.2 32.6 29.9 30.2 22.3 21.2 25.3 25.0 23.9 27.8 23.9 29.0 28.0 29.7 23.0 23.9 21.5 23.2 21.8 21.7 21.7 28.0 32.0 32.2 32.0 27.8 24.7 25.6 24.1 22.8 25.3 24.2 25.7 29.4 26.5 27.8 26.2 24.1 27.2 25.0 24.9 25.4 30.9 25.6 29.6 25.6 24.8 26.6 24.1 25.2 24.1 24.5 23.7 23.6 25.0 22.3 24.1 26.1 25.8 23.7 25.7 25.8 23.7 25.7 23.5 23.7 26.1

60.4 62.5 66.5 63.6 58.7 61.2 56.9 63.0 81.5 74.7 75.5 55.8 53.1 63.3 62.5 59.7 69.4 59.7 72.5 70.0 74.4 57.5 59.7 53.7 58.0 54.5 54.3 54.3 70.0 80.1 80.6 80.1 69.4 61.7 63.9 60.2 57.1 63.3 60.4 64.2 73.6 66.2 69.4 65.6 60.2 68.1 62.5 62.2 63.6 77.1 63.9 74.0 63.9 62.0 66.5 60.2 63.0 60.2 61.2 59.2 58.9 62.5 55.8 60.2 65.3 64.4 59.2 64.2 64.4 59.2 64.2 58.7 59.2 65.3 (continued on next page)

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Table 1 (continued) Latitude (degrees)

Longitude (degrees)

Curie point depths (km)

Geothermal gradient (°C/km)

Heat flow (mW m−2)

−14.5 −14.5 −14.5 −14.5 −14.5 −14.5 −15.5 −15.5 −15.5 −15.5 −15.5 −15.5 −15.5 −15.5 −15.5 −15.5 −15.5 −16.5 −16.5 −16.5 −16.5 −16.5 −16.5 −16.5 −16.5 −16.5 −16.5 −16.5

34.0 34.5 35.0 35.5 36.0 36.5 31.5 32.0 32.5 33.0 33.5 34.0 34.5 35.0 35.5 36.0 36.5 31.5 32.0 32.5 33.0 33.5 34.0 34.5 35.0 35.5 36.0 36.5

22.7 25.7 24.6 23.4 22.7 23.6 21.9 21.4 20.3 22.2 23.3 22.5 20.9 22.7 24.0 23.6 25.2 23.4 24.9 25.3 23.0 21.2 23.3 23.9 22.3 23.2 22.4 21.7

25.6 22.6 23.6 24.8 25.6 24.6 26.5 27.1 28.6 26.1 24.9 25.8 27.8 25.6 24.2 24.6 23.0 24.8 23.3 22.9 25.2 27.4 24.9 24.3 26.0 25.0 25.9 26.7

63.8 56.4 58.9 62.0 63.9 61.4 66.2 67.8 71.4 65.3 62.2 64.4 69.4 63.9 60.4 61.4 57.5 62.0 58.2 57.3 63.0 68.4 62.2 60.7 64.9 62.5 64.7 66.8

relates the CPD (Zb) to the heat flow (qz) is given by (Okubo et al., 1985; Turcotte and Schubert, 1982): qz ¼ −λ:½dT=dz ¼ −λ:½θc=Zb 

ð5Þ

where the heat flow qz is in mW m−2, and the geothermal gradient dT/ dz is in °C/km. λ is thermal conductivity in Wm−1 °C −1. The heat flow is calculated at each CPD (Zb) using Eq. (5) assuming a Curie temperature of 580 °C (Hussein et al., 2013; Leseane et al., 2015; Ross et al., 2006; Tanaka et al., 1999). The thermal conductivity (λ) is also a variable and depends on the lithology. Thus, proper knowledge of the lithology of the study area is required in order to obtain the representative result of the heat flow values. The pre-rift rocks in the MRZ are characterized by Precambrian crystalline metamorphic and igneous rocks and Paleozoic-Mesozoic sedimentary rocks (Carter and Bennett, 1973; Chilton and Smith-Carington, 1984). The Precambrian basement complex in the MRZ is dominated by gneisses, granulite, and schists (Bloomfield, 1966). Laboratory experiments on granulite-facies rocks suggest thermal conductivity values can range between 2 and 3 Wm−1 °C−1 (Ray et al., 2015). In addition the thermal conductivity ranges between 1.3 and 3.3 Wm−1 °C−1 for granites and 2.5–5.0 Wm−1 °C−1 for metamorphic rocks (Chilton and SmithCarington, 1984; Lillie, 1999). Given that the typical basement rock types in the MRZ are silicate rich igneous and metamorphic rocks, an average thermal conductivity value of 2.5 Wm−1 °C−1 is used in this study, which is both an average value for these types of rocks and a typical average crustal conductivity (Clauser and Huenges, 1995; Fagereng, 2013; Petitjean et al., 2006). 4. Results Table 1 shows the results of the CPD, geothermal gradient and the heat flow estimated from the aeromagnetic data. The CPD values beneath the MRZ range between 18 and 27 km (Fig. 5) and the geothermal gradients range between 22 and 32 °C/km (Fig. 6), while the heat flow values range between 59 and 82 mW m−2 (Fig. 7). Figs. 5–7 respectively show a well-defined NE trending zone of shallower CPDs (18–20 km),

high geothermal gradient (28–32 °C/km), and high heat flow (70–82 mW m−2) that extend beneath the center of the Malawi Rift, to the Ruhuhu and Maniamba troughs. This NE trending region of high heat flow (Fig. 7) corresponds to the region with the thinnest crust determined from passive seismic receiver functions (Fig. 6; Borrego et al., 2018) with crustal thickness ranging between 30.6 and 38 km. There is a localized region of high heat flow (70–74 mW m−2) beneath the RVP with CPD ranging between 20 and 22 km and a slightly elevated geothermal gradient (26–28 °C/km) (Figs. 5–7). Based on passive seismic receiver function analysis, the average crustal thickness beneath the RVP is 39 km (Fig. 6; Borrego et al., 2018) while Njinju (2016) used power-spectral analysis of gravity data to determine an average crustal thickness of ~43 km beneath the RVP and suggested the presence of an underplated magmatic body whereby, the gravity-derived crustal thickness (~43 km) represent the base of the underplated magmatic body whereas the seismically-derived crustal thickness (39 km) represent the top of the underplated magmatic body. Further south, there is another E-W trending region of shallow CPD (18–21 km), slightly elevated geothermal gradient (26–28 °C/km) and high heat flow (70–74 mW m−2) that extends beneath the Shire Graben and the Zambezi Rift (Figs. 5–7). The analysis of Rayleigh waves by Wang et al. (2019) showed that this region has a crustal thickness of 36–38 km beneath the Shire Graben. Apart from the thermal anomaly beneath the RVP, regions of high heat flow are coincident with the Karoo rift basins. Figs. 5–7 also show two regions of deep CPDs (24–27 km), low geothermal gradient (22–24 °C/km), and low heat flow (53–63 mW m−2). The first region is an E-W trending area that extends beneath the Bangweulu cratonic block to the western flank of the northern segment of the Malawi Rift where seismically derived crustal thickness measurement produced maximum values of 50 km (Fig. 6; Borrego et al., 2018). The second region is an E-W trending area that extends beneath the Niassa Craton through the Monkey Bay segment of southern Malawi Rift (Figs. 2c and 8). The low heat flow values are associated with thicker crust (38–40 km) under the Niassa Craton (Wang et al., 2019).

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Fig. 5. Map of the Curie Point Depth (CPD) values beneath the Malawi Rifted Zone obtained from two-dimensional (2D) radially averaged power spectrum analysis of the aeromagnetic data in Fig. 3. The CPD map is draped onto Shuttle Radar Topography Mission (SRTM) Digital Elevation Model (DEM). Dotted brown lines represent shear zones (SZ). The location of the Niassa craton enclosed in a black line is modified after Westerhof et al. (2008). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

5. Discussions 5.1. Thermal structure beneath the Malawi Rifted Zone Overall, the range of heat flow values for the MRZ are in good agreement with those reported from the Okavango Rift Zone and other segments of the magma-poor Western Branch of the EARS such as the Kivu Rift and Tanganyika Rift (Fig. 2a) where heat flow values range between 53 and 83 mW m−2 (Fadaie and Ranalli, 1990; Leseane et al., 2015). Higher heat flow values (relative to cratonic provinces) are typical in many active continental rifts (Olsen, 1995; Thybo and Nielsen, 2009). Even one of the world's youngest continental rifts (Okavango Rift Zone in Botswana) shows an elevated heat flow (N55 mW m−2)

(Leseane et al., 2015) at locations coincident with 4–5 km of elevated Moho along the rift axis (Fadel et al., 2018; Yu et al., 2015). In the MRZ, the distribution of CPDs and associated heat flow values show that not all of the rift basins have elevated heat flow (Figs. 5–7). Intriguingly, the regions of highest heat flow (70–82 mW m−2) in the MRZ predominantly concentrate within the Permo-Triassic Karoo rift basins including the Ruhuhu Trough, Maniamba Trough, the Shire Graben and the Zambezi Rift. The active Neogene Malawi Rift shows no simple north-south gradient in elevated heat flow along its axis (Fig. 7). Pronounced thermal anomalies only occur at the extreme north within the RVP and the central segment of the rift where the Mwembeshi and Macaloge Shear Zones obliquely intersect the Malawi Rift (Fig. 2c). The heat flow distribution along the Malawi Rift (Fig. 8a–

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Fig. 6. Map of the geothermal gradient beneath the Malawi Rifted Zone calculated from the Curie Point Depth (CPD) values, which are obtained from the two-dimensional (2D) radially averaged power spectrum analysis of the aeromagnetic data in Fig. 3. The geothermal gradient map is draped onto Shuttle Radar Topography Mission (SRTM) Digital Elevation Model (DEM). Dotted brown lines represent shear zones (SZ). The location of the Niassa craton enclosed in a black line is modified after Westerhof et al. (2008). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

b) is consistent with earlier observations from lake bed measurements (Fig. 8c; von Herzen and Vacquier, 1967) and onshore hot spring temperatures (Fig. 8d; Atekwana et al., 2015; Gondwe et al., 2012; Kinabo, 2016). Although heat flow from lake bed measurements at the center of the lake have a maximum value of 120 mW m−2 (Fig. 8c; von Herzen and Vacquier, 1967), studies by Fadaie and Ranalli (1990) suggest a mean heat flow value of 82 ± 78 mW m−2 after being corrected for lake sediment blanketing and lake circulation. It is well known that upper crustal magmatism provides convective heat transport to the earth's surface by the discharge of hot fluids such

as magma and hydrothermal fluids (e.g., Lysak, 1992; Kearey et al., 2013). This explains the coincidence of high heat flow with the RVP, which is the only region in the MRZ with known Neogene volcanism (Ebinger, 1989; Ebinger et al., 1989; Fontijn et al., 2010, 2012). Recent seismic studies imaged a pronounced low velocity zone beneath the RVP (Accardo et al., 2017; Grijalva et al., 2018) which further support the magmatic origin of the elevated heat flow beneath the RVP. Overall, the absence of consistent pattern of elevated heat flow along the Malawi Rift suggests that the crust has not been stretched to a sufficient extent (Wang et al., 2019) to allow for magma to ascend to shallow crustal

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Fig. 7. Map of the terrestrial heat flow beneath the Malawi Rifted Zone calculated from the Curie Point Depth (CPD) values, which are obtained from the two-dimensional (2D) radially averaged power spectrum analysis of the aeromagnetic data in Fig. 6. The heat flow map is draped onto Shuttle Radar Topography Mission (SRTM) Digital Elevation Model (DEM). Dotted brown lines represent shear zones (SZ). SHmax orientation from Heidbach et al. (2016). The location of the Niassa Craton enclosed in a black line is modified after Westerhof et al. (2008). MG = Malombe Graben, MJT = Makanjira Trough, SG = Shire Graben, ZG = Zomba Graben. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

level and most of the extension due to the Neogene rifting in Malawi is most likely preferentially localized in the lithospheric mantle. However, the presence of relatively thinner crust (30–38 km) beneath the Ruhuhu and Maniamba troughs in the eastern flank of the northern segment of the Malawi Rift compared to thicker crust (46–50 km) in the western flank of the rift (Fig. 6; Borrego et al., 2018), suggests that crustal thinning resulting from extension during the Permo-Triassic rifting in the Ruhuhu and Maniamba troughs still persist. In the central Malawi Rift, there is no record of upper crustal magmatism, yet heat flow values reach a maximum of 70–82 mW m−2 and hot spring temperatures attain a regional high of 79.3 °C (Fig. 8b– d). Geochemical analyses of gases from hot springs from the central Malawi Rift suggest the lack of mantle contribution (Atekwana et al., 2015). However, this segment of the Malawi Rift overprints the

southwestern terminations of Permo-Triassic Karoo rift basins (Ruhuhu and Maniamba Troughs; Figs. 2 and 8a), which ubiquitously show the highest heat flow in the MRZ. We interpret this anomalous heat flow in the Karoo rift basins is related to persistent residual heat flow originating from the Permo-Triassic rifting events. Based on previous global surface heat flow studies (Chapman and Pollack, 1975; Polyak and Smirnov, 1968), temporal dissipation of heat flow through continental provinces of Karoo age is about 58 ± 12 mW m−2, and at least 73 ± 10 mW m−2 through provinces younger than Late Eocene. This suggests that our estimated heat flow in the MRZ Karoo rift basins considerably exceeds the global mean residual heat flow through continental tectonic provinces of similar age (Figs. 8b–c). Segments of the Karoo rift basins show moderate crustal thinning (~35 km; Borrego et al., 2018; Njinju, 2016) and underplated magmatic

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bodies along the rift segments (Matende, 2015; Njinju, 2016). However, the post-Karoo (Cretaceous) emplacement of igneous intrusions in the MRZ (Fig. 2b; Castaing, 1991) provides a mechanism for the thermal replenishment of the initial decaying heat flow associated with the Karoo rifting. The Cretaceous dike swarms propagated beyond their source and extend for N280 km towards other Karoo rifts with the emplacement of ring-shaped igneous complexes at off-rift locations (e.g., Fig. 2b; Castaing, 1991; Heilman et al., 2019). Additionally, magnetotelluric imaging of the Luangwa Rift (Sarafian et al., 2018) provides geophysical evidence for possible fluid alteration of the Mwembeshi Shear Zone at a location, which coincides with high heat flow in our study (Figs. 7–9). This suggests the possible transport of hot fluids through Karoo rift basins that reactivated the Precambrian shear zones. In addition, current NNE-striking SHmax at the location (Fig. 7; Heidbach et al., 2016) supports the potential for fracture permeability to hydrothermal circulation. The presence of uranium enrichment in the Karoo basins may also provide additional source of heat flow. The presence of uranium mineralization in these rift basins is evidenced by the ongoing mining of uranium deposits hosted by Karoo sedimentary rocks in the Ruhuhu Trough (Ruhuhu and Mtonya Uranium Projects), in south Tanzania (Fig. 7; Hankel, 1987; Kögler et al., 1983; Langer et al., 2017). The presence of heat producing sources in the crust might explain the origin of the hot springs near the Keyalikera Uranium Mine (Becker et al., 2014; Fig. 7) in northern Malawi, a location we find to have relatively lower surface heat flow (Fig. 7). Therefore, we suggest that the elevated heat flow in the Karoo rift basins may have been promoted by the postKaroo diking events, possible hydrothermal fluid circulation and uranium enrichment. A potential analogue for uranium enrichment in the intra-continental Karoo rift basins is the recently observed uranium enrichment in Jurassic rifts linked to Karoo rifting at the boundary between West and East Antarctica adjacent to the Weddell Sea Rift System (Leat et al., 2018). Two distinct regions of deep CPDs (24–27 km), low geothermal gradient (22–24 °C/km), and low heat flow (53–63 mW m−2) within the MRZ (Figs. 5–7) coincide with the Bangweulu cratonic block in the northwest, and the Niassa Craton to the south. These observations are consistent with known thermal signatures of Precambrian cratonic provinces (Kearey et al., 2013; Lysak, 1987; Nyblade et al., 1990; Nyblade, 1997; Sclater et al., 1980). Interestingly, we found that the eastern extent of the Niassa Craton has the lowest thermal signature and coincide with a ~195 km long rift gap where hot springs are largely absent along the Malawi Rift (Fig. 8d). In this rift segment (Monkey Bay), the rift bifurcates around the Shire Horst (Figs. 2b and 8a), into the Malombe Graben (MG) in the east and the Makanjira Trough (MJT) to the west. South of the Shire Horst, in the Zomba Graben (ZG), the rift splays into a zone of diffused faulting as it transitions into the NW-trending Shire Graben (Fig. 8a). Overall, the southward shallowing of the Malawi Rift (Ebinger et al., 1987; Laó-Dávila et al., 2015) and its bifurcation and diffused faulting at its intersection with the relatively ‘cold’ Niassa Craton in the south demonstrates the influence of buried microcratons on the development and propagation of juvenile magma-poor rift segments. Likewise, the termination of the Luangwa Rift within the Bangweulu Block may represent a similar process. 5.2. Terrestrial heat flow distribution in active magma-poor rift zones It is well established that surface heat flow in any continental province is strongly related to the age of the province (e.g., Lee and Uyeda,

13

1965; Pollack, 1982; Polyak and Smirnov, 1968; Sclater and Francheteau, 1970). Surface heat flow generally diminishes with increasing time since the last tectonic event, with the presently active basins exhibiting the greatest heat flow (e.g., Kearey et al., 2013). Consequently, heat flow along continental rifts generally decreases with increasing time since the last rifting event, with the presently active rift segments exhibiting the greatest heat flow. Our estimated heat flow distribution in the MRZ suggests that this simple rule may be consistent in areas of active magma-assisted rifting, but will likely not apply to active magma-poor rift zones such as the MRZ. It is reasonable to find that magma-rich segments of continental rift systems have higher surface heat flow compared to the magmapoor segments of similar age, evident in surface heat flow distribution along the active western and eastern branches of the EARS (Lysak, 1987). Along the magma-poor western branch, elevated heat flow generally concentrates at the few volcanic provinces occurring along the rift branch. Based on our findings, we propose that in active magma-poor rift zones, pronounced heat flow will concentrate not only in isolated zones of magmatic activity, but may also focus largely within nearby ancient magmatic and/or thinned-crust rift segments (Fig. 9). This observation suggests that the great variability of heat flow in continental rifts (Fig. 1b) may be explained by contributions from tectono-thermal inheritance, thus posing important implications for active continental rift zones in other places (Fig. 1a). Further, we propose that Karoo rift basins are areas of potential geothermal energy exploration in eastern Africa. 6. Conclusions In this study we determined the thermal structure beneath the Malawi Rifted Zone (MRZ) using the 2D radially averaged power spectrum analysis of aeromagnetic data. We found a shallow CPD (18–20 km), high geothermal gradient (29–32 °C/km), and elevated heat flow (70–82 mW m−2) beneath the Permo-Triassic Karoo rift basins in the MRZ. We do not find any consistent pattern of elevated heat flow beneath the axis of the active Neogene Malawi Rift except for localized regions of elevated heat flow (70–82 mW m−2) beneath the RVP and the central segment, where the Malawi Rift overprints Karoo Ruhuhu and Maniamba Troughs. The absence of a consistent pattern of elevated heat flow along the active Malawi Rift axis suggests that the crust has not been stretched to a sufficient extent and most of the extension due to the Neogene rifting in Malawi is most likely preferentially localized in the lithospheric mantle. Further, our findings suggest that in active magma-poor rift zones, pronounced heat flow will concentrate not only in isolated zones of magmatic activity, but may also focus largely within nearby ancient magmatic and/or thinned-crust rift segments. Thus, our findings demonstrate the role of tectonothermal inheritance in the thermal subsidence pattern of continental rift basins. Finally, we suggest that important targets for geothermal exploration in the MRZ include the thermal anomaly beneath the RVP from Cenozoic magmatism and residual heat flow in the Karoo basins due to Mesozoic rifting. Acknowledgments This project was supported by the National Science Foundation – Continental Dynamics (NSF-CD) grant # EAR-1255233 and Office of International and Integrative Activities grant # IIA-1358150. The aeromagnetic data were provided to us from the Council of Geosciences, South Africa, as part of the Southern African Development Community

Fig. 8. (a) Shuttle Radar Topography Mission (SRTM 30 m resolution) Digital Elevation Model (DEM) showing the Malawi Rift. (b) Along-rift-axis heat flow values from this study, taken at locations shown as stars (x) in Fig. 8a. Red dotted lines represent global mean heat flow through continental tectonic provinces of Mid-Permian and Late-Eocene ages (after Polyak and Smirnov, 1968; Chapman and Pollack, 1975). (c) Along-rift-axis lake bed heat flow measurements (von Herzen and Vacquier, 1967), obtained at locations shown as circles (o) in Fig. 8a. (d). Along-rift onshore hot spring temperatures from Atekwana et al. (2015), Gondwe et al. (2012) and Kinabo (2016) obtained at locations shown as triangles in Fig. 8a. The horizontal axes of Figs. 8b-d are latitudes north (on the left) to south (on the right). MT = Maniamba Trough, MG = Malombe Graben, MJT = Makanjira Trough, RT = Ruhuhu Trough, SG = Shire Graben, SH = Shire Horst, UT = Usangu Trough, ZG = Zomba Graben. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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Fig. 9. Three-dimensional (3D) conceptual model of the magma-poor Malawi Rifted Zone (MRZ) illustrating how the residual heat flow in regions of ancient (Mesozoic) rifting dominates the landscape. The active Malawi Rift (Cenozoic) generally has low heat flow except in an isolated area of active magmatism in extreme north (Rungwe Volcanic Province), and the central segment where the rift overprints Karoo rifts (Ruhuhu Trough and Maniamba Trough). The coincidence of electrical conductivity of the Precambrian Mwembeshi Shear Zone (MSZ) and location of elevated heat flow in the Luangwa Rift may suggest processes associated with hydrothermal fluid circulation through fracture systems that reactivated the shear zone. Current NE-striking SHmax at the location also suggests the potential for fracture permeability to hydrothermal circulation.

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